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Chapter 16
The Marine Cryosphere
David M. Holland
Courant Institute of Mathematical Sciences, New York University, New York, USA
Chapter Outline
1. Introduction
1.1. Marine Cryosphere
1.2. Ice Physics
1.3. Ocean Impacts
1.4. Relation to Other Chapters
2. Sea Ice
2.1. Observations
2.2. Modeling
2.3. Ocean Mixed-Layer Interaction
2.4. Polynyas
2.5. Impact on Water Masses, and Circulation
2.6. Biogeochemical Ramifications
3. Land Ice
3.1. Observations
3.2. Modeling
3.3. Ocean Mixed-Layer Interaction
3.4. Impacts on Water Masses
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1. INTRODUCTION
1.1. Marine Cryosphere
The cryosphere is loosely defined as the component of the
earth’s climate system consisting of frozen water. In one
way or another, all components of the cryosphere interact
with and influence the ocean. The cryosphere can be considered to consist of sea ice, land ice, and atmospheric
ice. While atmospheric ice, essentially consisting of frozen
water clouds, has an indirect impact on the ocean through
precipitation, including snowfall onto sea ice and land
ice, and atmospheric radiation influences, we discuss it
no further here as it is more of an atmospheric science phenomenon. We focus on ice that is more or less in direct
contact with the ocean (see Figure 16.1). Land ice, often
referred to as a glacier, is in direct contact with the ocean
wherever it flows, under the action of gravity, into the ocean
to form floating ice shelves along the periphery of ice sheets
and ice caps. A particular type of land ice, marine
Ocean Circulation and Climate, Vol. 103. http://dx.doi.org/10.1016/B978-0-12-391851-2.00016-7
Copyright © 2013 Elsevier Ltd. All rights reserved.
3.5. Geochemical Tracers
3.6. Sea-Level Change
4. Marine Permafrost
4.1. Pure Ice
4.2. Methane Clathrates
5. Emerging Capabilities
5.1. Ice-Capable Observations
5.2. Ocean-Capable Observations
5.3. Ice-Capable Modeling
6. Cryospheric Change
6.1. Observed Sea-Ice Change
6.2. Sea-Ice Projections
6.3. Observed Land-Ice Change
6.4. Land-Ice Projections
6.5. Marine Permafrost
7. Summary
References
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permafrost, is also found beneath the seafloor, where it
holds a potential for interaction with the ocean. Sea ice,
by its very nature, is formed at the surface of the ocean
and is perhaps the most well understood and studied aspect
of ice–ocean interaction. Depending upon the time of year,
up to 5% of the global ocean surface is covered by ice
(NSIDC, 2007). Collectively, we refer to the sea ice,
glacier, and permafrost elements as the marine cryosphere.
The interactions of sea ice, land ice, and marine permafrost with the oceans are largely confined to the polar
oceans, with the only exception being some marine terminating glaciers found south of the Arctic Circle in, for
instance, Alaska (Meier et al., 1980), and north of the
Antarctic Circle in Chile (Warren, 1993). The bathymetric
layouts of the far northern and southern oceanographic
basins are essentially morphological opposites. Looking
to the north, the Arctic basin is largely landlocked, the only
exchanges with the rest of the global ocean occurring with
the North Atlantic Ocean through the Fram Strait and
Barents Sea passages, as well as the modest connectivity
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(b)
FIGURE 16.1 Photographs of typical (a) sea ice and (b) iceberg and ice-shelf areas of the polar oceans. Sea ice is marine in origin originating from
freezing of sea water. Icebergs are meteoric ice, born from ice shelves, originating with evaporation of sea water that falls as snow onto the great ice
sheets. http://nsidc.org/cryosphere/quickfacts/seaice.html; http://summitvoice.files.wordpress.com/2011/07/ice.jpg.
to the Pacific Ocean through Bering Strait, and some
exchange through the Canadian Arctic Archipelago. This
geographic restriction makes it somewhat easier to observe
the water mass exchanges in the north as compared with the
south (see Chapter 17). The bathymetry of the Southern
Ocean is completely open to the global ocean, with no
restriction to exchange with the neighboring South Atlantic,
South Pacific, and Indian Oceans (see Chapter 18).
The difference in land configuration in the north and
south is a key factor in the distinction between the sea-ice
and land-ice cover of the two regions, discussed at length
in Sections 2 and 3, respectively. The Arctic Ocean, consisting principally of the area north of 80 N, is surrounded
by land that is largely nonglaciated and has river runoff at
the surface in the boreal summer that remains largely trapped
in the Arctic basin leading to a relatively stratified ocean and
thick sea-ice cover. The Southern Ocean, essentially in an
unconfined basin and with almost no surface runoff of freshwater, is more weakly stratified and has a much thinner seaice cover compared with the Arctic. In other words, there is
not really an “Antarctic Ocean” analog to the Arctic Ocean;
instead, there exists the Southern Ocean, much of which is
subpolar and cannot be compared in any real sense to the
truly polar Arctic Ocean.
There are a number of distinctions to be drawn between
the coastal and the continental shelf zones of the two polar
regions. The Arctic continental shelves cover approximately 10% of the area of the Arctic Ocean and have an
average depth of little more than 100 m (Jakobsson,
2002). The situation in the south is fundamentally different
with the continental shelf around Antarctica occupying a
much smaller percentage area of the Southern Ocean and
the average depth of the continental shelf being far greater,
on average 600 m (Anderson et al., 1980). The global mean
continental shelf depth is estimated to be 150 m (Bott,
1971), and thus, the Antarctic shelf is often referred to as
being overdeepened. The deepness is a result of cumulative,
erosive activities of a grounded ice sheet (SCAR, 1997).
This distinction in average depth is of fundamental importance to the manner in which warm subpolar waters can or
cannot go aboard the continental shelves and reach the
periphery of the major ice sheets. Another distinct feature
of the Antarctic continental shelf is that it is foredeepened,
that is, the shelf depth becomes deeper as one travels from
the shelf break in toward the margin of the ice sheet (Pope
and Anderson, 1992).
The modern marine cryosphere of the north consists of
perennial Arctic sea-ice cover, the Greenland Ice Sheet, and
marine permafrost; in the south, there is the seasonal sea-ice
cover and the more massive Antarctic Ice Sheet. On longer
timescales, back through the Pleistocene, the northern cryosphere has undergone larger areal changes as ice sheets
occupied the Canadian (Broecker et al., 1989) and Scandinavian (Rinterknecht et al., 2006) land masses more often
than not during that epoch.
The character of the seasonal and perennial sea-ice
cover of the polar seas has been well noted by voyagers
dating back over the past millennium, certainly as early
as the time of the Vikings in the north (Ogilvie et al.,
2000). In more recent times, over the past century or so,
explorers and oceanographers have set sail deep into both
the Arctic Ocean (Nansen, 1902) and the Southern Ocean
to the coast of the Antarctic continent (Ross, 1847). The
mere existence of sea ice, aside from its physical implications for ocean circulation in the high latitudes, has in itself
greatly hampered the exploration and understanding of
polar ocean dynamics. The difficulty of cutting through a
meter or more of frozen ocean surface and the associated
risks of being trapped in the “pack” for a year or more have
brought many an expedition to a tragic end (e.g., the
Franklin expedition; Berton, 1988) or at least a difficult
ending (e.g., Amundsen’s travel through the Northwest
Passage; Berton, 1988). In the 1890s, Nansen deliberately
set his ship into the Arctic ice pack on the Siberian coast,
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and riding with the transpolar drift, he crossed close to the
North Pole, ultimately to emerge into the Fram Strait area,
thereby demonstrating the basic circulation pattern of
Arctic sea ice (Nansen, 1902). In a somewhat analogous,
albeit not so planned adventure, Ernest Shackleton was
set into the pack of the Weddell Sea, and through the trajectory of his ship and travels, the circulation pattern of
the ice of the Weddell Sea was better established
(Shackleton, 1919).
While sea ice is an evident participant in the marine cryosphere, perhaps not so obvious is the role of land ice. Both the
continent of Antarctica (Drewry et al., 1983) and the island of
Greenland (Bamber et al., 2001) are almost everywhere
covered by an ice sheet several kilometers in thickness. In
some locations, these ice sheets actually sit upon bedrock
that is below modern-day sea level. Such areas, referred to
as marine ice sheets (Mercer, 1978), have direct interaction
with the ocean. In some instances, the marine ice sheets terminate at the ocean in vast extensive floating ice shelves,
such as in the Ross (Shabtaie and Bentley, 1987) and
Weddell seas (Swithinbank et al., 1988), while in other cases,
the termination can be rather more abrupt with the glacier
only presenting a rather modest-sized interface to the ocean,
as in the case of many outlet glaciers in Greenland (Csathó
et al., 1999). Large portions of the West Antarctic Ice Sheet
were discovered to be well below sea level during expeditions undertaken in the International Geophysical Year
(Bentley and Ostenso, 1961). Such marine-based glaciated
areas can be thought of as being an extension of the global
ocean. There is some evidence from seafloor sediment
records that the ocean has occupied these areas from time
to time during past interglacial periods (Scherer et al.,
1998). It is the waxing and waning of the marine ice sheets
that holds the greatest present interest for potential rapid
global sea-level change in the current century and beyond
(Alley et al., 2005).
1.2. Ice Physics
To better understand the manner in which ice interacts with
the ocean, we first review a few of the fundamental physical
properties of ice and freezing ocean water. The water molecule, H2O, has more solid phases than any other known
substance. The most common form of solid ice takes on a
hexagonal pattern, termed hexagonal ice, or ice Ih, or
simply ice. Such ice, created from water at temperatures
below 0 C, settles into a regular crystalline structure with
hexagonal symmetry (Hobbs, 1971). At even lower temperatures, ice crystallizes into structures with cubic symmetry,
but we are only concerned in an oceanographic context with
the hexagonal form. One exception is clathrates, discussed
in Section 4 in the context of marine permafrost.
A phase diagram for water, showing the occurrence of the
liquid, solid, and gas phases as functions of ambient
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temperature and pressure, reveals that the melting line has a
negative slope (Figure 16.2). This means that as the pressure
on ice is increased it tends to move from the solid toward the
liquid phase. This is a consequence of ice expanding as it solidifies, giving rise to the remarkable fact that the water solid
phase (Ih) is less dense than its liquid phase, so that ice shelves,
icebergs, and sea ice all float upon the ocean, rather than
resting on the seafloor, as would be the case if the ocean were
composed of almost any other liquid substance.
The mechanical strength of ice varies depending upon
the scale under consideration. At the smallest scale, that
is of an individual floe of sea ice or a small iceberg, ice
is a rigid crystalline solid, gaining its strength from
hydrogen bonding, and it fractures in a brittle fashion
(Petrovic, 2003). By contrast, on the large scale, sea ice
(Mellor, 1983) and ice sheets (Weertman, 1983) follow distinctly different flow regimes than at the small scale. We
will further investigate these large-scale dynamical regimes
in Sections 4.2 and 4.3 as they are of great consequence to
understanding the impact of ice upon the ocean. Deeper
insight into the small- versus larger-scale behavior of ice
comes from recognizing that, at the level of the crystal
structure, individual molecules are linked together in a
single crystal, but at the larger scale, ice is a polycrystalline
material and its strength is dependent on that structure
(Hooke, 1981). Ice has both an elastic behavior, arising
from small displacements from their equilibrium position
of atoms in a crystal, and a plastic behavior which emerges
from the motion of dislocations that are the defects in the
otherwise perfect crystal structure. An in-depth treatment
of the molecular and aggregate strength properties of ice
is found in the textbook by Petrenko and Whitworth (1998).
The thermodynamical properties of ice govern its
growth and decay. One has to consider the density, specific
heat, latent heat, and related thermodynamical properties of
ice to model its evolution in a given thermodynamic regime.
For sea ice, which is marine in origin, the thermodynamics
can be complicated by the presence of significant amounts
of brine locked into the ice (Cox and Weeks, 1982), something not found in glacier ice which is generally meteoric in
origin. Whereas freshwater reaches maximum density at a
temperature of approximately 4 C, the impact of salinity
in typical seawater is to move the temperature of maximum
density to the freezing point (Fujino et al., 1974). A
thorough review of ice thermodynamics is provided in
the textbook by Lock (1990), which delves into the thermodynamical Gibbs function, which allows one to develop the
phase diagram for ice, and thus to understand the manner in
which water switches between its phases of liquid, gas, and
solid (see also Chapter 6).
The freezing point of water is dependent upon both
pressure and salinity (Fujino et al., 1974). The details of
these dependencies have a fundamental impact on ice formation and decay in the ocean. The near-ubiquitous
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FIGURE 16.2 Phase diagram for water shown logarithmic in pressure and linear in temperature. The roman numerals indicate various ice phases. The
most common form of ice, Ih, and that most relevant to understanding ice–ocean interaction are located near the center of the diagram. http://en.wikipedia.
org/wiki/Phase_diagram.
presence of salt in the ocean lowers the freezing point below
the nominal 0 C mark of pure water, reaching down to
2 C for typical ocean salinity. The deeper a piece of
ice is found in the ocean, the greater the pressure and the
lower the freezing point yet again. Combined, the effects
of salinity and pressure can reduce the freezing point for
some deep glaciers to below 3 C. The effect of the depth
dependence on the freezing temperature can result in more
rapid melting for the deeper parts of a glacier than for its
shallower parts.
Salinity plays a crucial role in the melting of ice for yet
another reason. The molecular diffusivity of salt is approximately one hundred times smaller than that of heat, and as a
result, the transfer of heat at the interface where ice and
water meet is much more efficient than that of salt. This
is because turbulence is suppressed in the upper few millimeters of the boundary layer so that exchanges between ice
and ocean occur only by molecular diffusion (Josberger and
Martin, 1981). This diffusivity contrast has the effect of
making the interface less salty than the ambient fluid during
melting, leading to rising of the freezing point, and ultimately slowing of the rate of ice melt compared to the rate
that would occur if such an effect were absent (Holland and
Jenkins, 1999). In the case of freezing, the situation is
further complicated by the presence of convection at the
interface due to the unstable fluid stratification that occurs
when ocean water freezes and salt is rejected, by and large,
from the newly forming ice.
1.3. Ocean Impacts
The presence of an ice cover, whether sea ice or ice shelf,
has a profound effect on the underlying upper ocean.
Perhaps most notable is the transformation of a low-albedo
surface (typical ocean value of 0.10) to a high-albedo
surface (typical snow cover surface value of 0.90). The
abrupt decrease in transmission and increase in reflection
of shortwave radiation as an open-ocean area is converted
to a sea-ice-covered area completely alter the surface
energy balance (Allison et al., 1993). This gives rise to
the so-called ice-albedo effect whereby an increase in
sea-ice cover leads to less surface absorption of heat energy,
which in turn promotes further growth of sea ice and further
decrease in surface energy absorption (Robock, 1980). In a
paleoclimate context, such an effect has been widely implicated in the “snowball earth” phenomenon (Hoffman and
Schrag, 2002) whereby almost the entire global ocean
surface is covered by ice. The white surfaced planet earth
absorbs little surface radiation in such a scenario, thereby
locking the planet into an icy state. The converse, a planet
with no ice cover absorbs an enormous amount of radiation,
allowing it to stay in an ice-free state. Likely, a bifurcation
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point exists in the earth’s climate system allowing the earth
to toggle between ice-covered and near-ice-free states, like
the present (North et al., 1981).
Aside from its impact as a radiative barrier, the sea-ice
cover is a highly effective barrier to gas exchange between
ice and ocean. A key greenhouse gas, carbon dioxide, is
continuously exchanged between the ocean and atmosphere, the rate of exchange being controlled by the partial
pressure of gas in the atmosphere and the wind speed, and
the temperature and salinity of the ocean, which controls the
solubility of the gas in seawater (Wanninkhof, 1992) (see
also Chapter 30). It is known from observational data that
this exchange is highly seasonal dependent, demonstrating
the key moderating role played of the ice cover (Golubev
et al., 2006). In ice-shelf-covered portions of the ocean,
there is no atmospheric-ocean exchange of gas, but there
is input of gas into the ocean where the base of an ice shelf
melts, as gases dissolved in the ice are released into the
ocean, providing a glacial source of dissolved oxygen
(Rodehacke et al., 2007).
Throughout the global ocean, the cycle of surface evaporation and precipitation, driven by the atmospheric state
and ocean surface conditions, ice-covered or otherwise,
helps set the salinity of the ocean surface. Salinity is the
determining factor on ocean density in cold waters in part
because the thermal expansion coefficient of water
decreases with decreasing temperature, hence the importance of surface freshwater fluxes at high latitudes in controlling the downwelling branches of the global
thermohaline circulation (THC), also known as the global
conveyor belt (Manabe and Stouffer, 1995). The formation
of sea ice at high latitudes results in the injection of salt into
the surface ocean, often resulting in unstable stratification
leading to deep or bottom convection (Aagaard and
Carmack, 1989), enhancing the conveyor belt. On the flip
side, the melting of sea ice along the sea-ice margin can lead
to an input of freshwater that can suppress or even shut
down convection, thereby slowing down the conveyor-belt
transport of heat from low latitudes to high latitudes.
Particular to the Arctic is the input of substantial river
runoff mentioned earlier ( 3500 km3/year, Lammers
et al., 2001); in fact, some 10% of the global river runoff
goes into the Arctic basin (Peterson et al., 2002). The
arrangement of basin layout and river input in the Arctic
leads to the strongest halocline to be found anywhere in
the global ocean (Nguyen et al., 2012).
The strong halocline effectively insulates the Arctic seaice cover from the vast, latent heat supply of the Atlantic
layer below and thereby allows a substantial thickness, at
least historically, of sea-ice cover to persist in the Arctic
Ocean (Steele and Boyd, 1998). By contrast, the Southern
Ocean has almost no surface freshwater runoff from the
Antarctic continent, and the water column is at best marginally stable and thus more easily allows for the upwelling
of deep, warm water (Foldvik and Gammelsrd, 1988) that
can contribute to a thinner sea-ice cover. The melting at the
base of the ice shelves and calving of icebergs cannot really
be thought of as a kind of “runoff,” as the melting related to
those processes ( 2400 km3/year; Rignot et al., 2013)
occurs largely at depth within the ocean. The resulting,
diluted meltwater contributions are evidently not large
enough in quantity to enforce substantive stratification in
the surrounding ocean. Massive, persistent convective
events, such as the Weddell Polynya event of the mid
1970s, are indicative of this borderline stability in the south
(Holland, 2001).
Throughout most of the global ocean, the input of turbulent kinetic energy into the surface ocean leads to a
downward mixing of surface properties over a depth scale
referred to as the mixed-layer depth (de Boyer Montegut
et al., 2004). This depth represents a balance between
buoyancy input at the surface ocean, which tends to thin
the mixed layer, and mechanical stirring, which tends to
deepen the layer. Over a sea-ice-covered ocean, as compared to the open ocean, this balance is altered both in terms
of buoyancy and mechanical input details, but the same fundamental principles apply (Lemke et al., 1990).
The mixed layer beneath an ice shelf is another distinct
environment for mixed-layer physics (Holland and Jenkins,
2001). Whereas sea ice does move with respect to the ocean
surface, an ice-shelf base does not, and consequently, there
is much less mechanical input into the mixed layer at the
base of an ice shelf. That notwithstanding, the base of an
ice shelf tends to be inclined somewhat and thus an effective
gravity current can flow upward along the basal slope,
leading to a unique source of mechanical mixing, different
from that beneath sea ice, which tends to be relatively flat
on such scales. A one-dimensional numerical model of a
plume flowing up the base of an ice shelf has been provided
by Jenkins (1991).
The final impact of ice on the ocean that we mention is
that of global sea-level change. Ice floating upon the ocean,
such as sea ice or ice shelves, makes only a minor contribution to sea-level change as the ice upon melting freshens
the relatively salty ocean, thereby raising the sea level
through modest mass contributions (Jenkins and Holland,
2007). More impressive, land ice grounded on bedrock
below current-day sea level has a volume above flotation
(i.e., the volume of grounded ice above and beyond that
which is simply displacing ocean water) that can contribute
substantially to global sea-level change. Marine-based West
Antarctica has been estimated to hold more than 3 m of
potential global sea-level change (Bamber et al., 2009).
1.4.
Relation to Other Chapters
This chapter, while dealing primarily with the physics and
phenomenology of sea ice and land ice, implicitly draws in
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material that is discussed elsewhere in this book, such as
mixed-layer dynamics in relation to sea-ice cover and bottom
water formation in relation to sea-ice growth and land-ice
basal melting. The state of sea ice and land ice in the polar
regions is greatly influenced by oceanic heat transport from
the subpolar oceans into the polar (see Chapter 11). The freshwater produced by melting sea ice and land ice may exert
influence on the overturning circulation in both the Nordic
Seas (see Chapter 17) and the southern boundary of the
Southern Ocean (see Chapter 18), particularly in the deepwater formation regions of the Ross and Weddell Seas.
The outline of the remainder of this chapter is as
follows. Sea ice is discussed in Section 2, followed by land
ice in Section 3, and a brief mention of marine permafrost in
Section 4. Recent advances in observations and modeling
capabilities both for sea ice and land ice, as well as of the
waters of the polar oceans, are treated in Section 5, which
covers emerging capabilities. Finally, a review of documented, recent changes in the marine cryosphere and an
outlook for the twenty-first century are given in Section 6.
2.
SEA ICE
The surfaces of the polar oceans have a net energy loss
through the course of an annual cycle. An important consequence is that the surface of the ocean freezes over, forming
sea ice, and as solid ice is less dense than liquid water, the
ice floats upon the ocean surface, turning it from dark blue
to bright white. With the swing of the seasons alternating
between the northern and the southern hemisphere, the
sea-ice cover of the Arctic and Antarctic grow and retreat
out of phase with one another, producing one of the largest
seasonal phenomena on the earth’s surface (see
Figure 16.3). The Arctic sea-ice cover varies from a wintertime maximum of 15 106 km2 to a summertime
minimum of 7, whereas the Antarctic swings between 20
and 4 (Gloersen et al., 1992).
The detailed process by which sea ice forms depends on
the state of the ocean’s surface, as a more turbulent ocean
will lead to a somewhat different ice cover than a calm
ocean. In either case, atmospheric cooling of the ocean
results in its surface being slightly supercooled in the first
instance, which in turn initiates nucleation of ice crystals
and frazil ice. The frazil ice platelets coalesce, leading to
a type of ice known as nilas in relatively calm seas and
to pancake ice in more turbulent seas, where distinct floes
are discernible (ASPECT, 2012). Floes can be pushed to
ride up upon one another, thus thickening the cover. Once
a sufficient cover exists, additional growth can occur at the
underside of the ice, whereby loss of heat by conduction
through the cover leads to formation of congelation ice
directly to the underside.
Growth of ice can also occur at the ice surface. This
happens when snow cover accumulates on the ice surface
Ocean Circulation and Water Masses
to the extent that the weight of the snow is sufficient to
depress the freeboard of the ice below sea level, causing
flooding of the ice surface and conversion of wet snow to
ice, the resulting ice being referred to as snow ice
(Jeffries et al., 1997).
Ice that forms from the sea, that is, ice of marine origin,
has a significant salt content, at least initially. The salinity
of the ocean is approximately 35 psu, and when frazil ice
forms, this salt is rejected. However, the rapid formation
of sea ice leads to some salt being trapped into pockets
within the larger sea-ice conglomerate, giving the ice a bulk
salinity of approximately 3–4 psu (Niedrauer and Martin,
1979). These distributed pockets become very concentrated
in salt and do not freeze. Over time, however, they tend to
drain down through the ice under the action of gravity. Melt
water washing through the ice during summertime as the
result of surface melt pond formation is another factor that
gradually reduces the salt content of sea ice (Fetterer and
Untersteiner, 1998). Glacier ice, by contrast, is generally
of meteoric origin and has very limited salt content, which
is largely aeolian in origin.
A fully developed ice cover is usually relatively mobile,
blown around by the winds and dragged by the ocean currents. Most of the Arctic and Antarctica sea-ice cover is
mobile, but there are areas near the coast where the ice is
stagnant and referred to as landfast ice (Mahoney et al.,
2007). The main characteristics of an ice cover are its
thickness and concentration. The thickness of Arctic sea
ice ranges from thin, that is, 0.1 m, to thick, that is, up to
6 m, whereas Antarctic sea ice tends to be relatively thin
overall, generally spanning the range 0.1–1 m. In some
locations, the sea ice can pile up substantially upon itself,
forming ridges that can reach up to 20 m in thickness.
The term “concentration” is used to describe the fractional
areal coverage of the ocean surface by sea ice and ranges
between 0% and 100%. While in the case of landfast ice
the coverage is full, 100%, the majority of the ice cover
has leads, or fracture zones, in which the ocean surface is
directly exposed to the polar atmosphere. Even though
the leads occupy only a few percent of the sea-ice region,
it is because of the great efficiency with which they can
exchange heat and moisture with the atmosphere that they
are an important aspect of the ice cover.
A major distinction between sea ice of the north and that
of the south is the age of the ice. Arctic ice can have a significant multiyear fraction, meaning that the ice has survived several seasons—largely the result of the enclosed
nature of the Arctic basin. By contrast, Antarctic ice is
almost exclusively first-year ice as each summer season
the ice cover is largely melted away. A similarity of Arctic
and Antarctic ice is that they both have marginal ice zones,
that is, transition zones between the interior, densely concentrated pack ice and the ice-free open ocean. The marginal ice zone migrates outward during winter and
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March
September
February
September
FIGURE 16.3 Seasonal cycle of sea-ice cover in both hemispheres. Top panels: Arctic winter maximum in March and summer minimum in September.
Bottom panels: Antarctic summer minimum in March and winter maximum in September. Blue areas are open ocean, gray areas are continental land mass,
and white is sea ice. http://nsidc.org/cryosphere/sotc/sea_ice.html.
retreats inward during summer, giving rise to an area of
coverage known as the seasonal ice zone.
2.1. Observations
The traditional observational strategy for sea ice has been to
make observations from ships traveling near or through the
pack. Much of this data is dependent upon the subjective
evaluations of an onboard ice observer, leading to some concerns over data quality. An effort has been made to standardize the observations, leading to a more quantitative
global data set (ASPECT, 2012). From a ship, estimates
of sea-ice concentration, often quoted in “egg-chart” fashion
as fractions of eight, and sea-ice thickness are feasible.
While a single ship-based cruise provides a relatively small
areal data sample, taken cumulatively over many cruises, the
global data set is of value to sea-ice researchers.
Complementing the ship-based data set is a submarine
observational data set, based primarily on upward looking
sonar that can give an estimate of ice areal coverage and
draft (Moritz and Wensnahan, 2011). The draft can be converted to ice thickness, subject to assumptions about the ice
density and snow cover. The data can be used to evaluate
decadal and multidecadal trends in sea-ice thickness,
spanning back to 1958 (Kwok and Rothrock, 2009). Presently, this type of data only exists for the Arctic region.
Another approach used to collect sea-ice data along a
track has been the development and deployment of buoys
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placed on the surface of the ice, both in the Arctic (IABP,
2012) and the Antarctic (IPAB, 2012). The original version
of these buoys collected surface air temperature and sealevel pressure data, and reported their position, all highly
valuable as input to weather forecasting models and subsequent atmospheric reanalysis products. The buoys provided
a very clear view of the circulation pattern of Arctic and
Antarctic sea ice and the relation of that circulation to the
sea-level pressure fields, and hence wind patterns. More
recent development in sea-ice buoys now has them placed
on the ice surface but with instrumentation extending all the
way through the ice into the top few meters of the ocean.
The buoys now have the ability to measure the temperature
through the sea ice as well as the ice thickness and the entire
mass balance of an ice floe. Such buoys have been
developed that can not only be deployed in ice but also float
on the ocean in the instance that the ice floe melts entirely
(SIMB, 2012).
The largest leap forward in sea-ice observations
occurred in the 1970s with the deployment of microwave
satellite remote sensing instrumentation (Gloersen et al.,
1992). The data provided the first global view of the entire
Arctic and Antarctic sea-ice fields, their seasonal variability, and their interannual trends. Data sets were limited
initially to primarily sea-ice concentration and used a
variety of techniques to convert brightness radiance to concentration, producing broadly similar data products. More
recently, developments using particle tracking, or image
correlation techniques, with data sets from visible or active
microwave sensors have allowed sea-ice velocity fields to
be assessed, giving an impressive view of the large gyres
that exist with the ice packs (Vesecky et al., 1988). Most
recently, laser altimetry has provided the first glimpses of
sea-ice thickness, based on measurements of the freeboard
of ice (Hvidegaard and Forsberg, 2002).
2.2.
Modeling
To understand the physics of large-scale sea-ice behavior, a
model is needed, theoretical in nature and numerical in
implementation. From a pragmatic view point, a numerical
model once developed can be used to make projections of
sea-ice behavior. Sea-ice physics, on the large scale, is
complex, even more so than the ocean in at least one
respect—while the continuum hypothesis upon which
ocean modeling is based has some reasonable foundation,
it is not the case for sea ice. Sea ice is composed of individual floes at scales ranging from 1 km down to 1 m,
making the description of the sea-ice fluid exceptionally
challenging. That notwithstanding, parameterizations of
sea ice allow one to describe in a model the aggregate
behavior of the underresolved smaller-scale features. Based
on comparisons of model simulations with observational
data, it can be claimed that there is indeed some accuracy
Ocean Circulation and Water Masses
and skill in the current generation of large-scale sea-ice
models (Fichefet et al., 2003). The early-generation seaice models were “stand-alone” in the sense that the atmospheric and oceanic boundary conditions were fed to the
sea-ice model equations as fixed boundary conditions
(Flato and Hibler, 1992) as opposed to the present generation where the boundary conditions are dynamical, coming
from atmosphere and ocean models to which the sea-ice
model is coupled. Subsequent development of coupling
strategies has now led to the point where sea-ice models
are fully coupled with ocean and atmospheric general circulation models used in climate-change simulation scenarios
(IPCC, 2007). We next discuss the major conceptual components of a sea-ice model: ice thermodynamics (growth/
decay) and dynamics (motion).
Sea-ice thermodynamical equations describe the flow of
heat through the sea-ice medium (Maykut and Untersteiner,
1971). As the aspect ratio of ice is such that ice is thin, the
flow direction of heat is largely vertical. Brine pockets
aside, the flow of heat is governed by a one-dimensional
(vertical) equation describing diffusion of heat by
molecular processes, aided by the fact that the thermal conductivity of sea ice is a well-established quantity (Ebert and
Curry, 1993). The addition of a snow cover requires only a
step change in the thermal conductivity to a smaller value
(as snow is more insulating than ice) (Grenfell, 1991).
The main boundary conditions are that of the flux of heat
applied at the base of the sea ice and at its upper surface.
Here we describe the surface fluxes and defer the basal
fluxes to the mixed-layer discussion below.
At the surface of the ice (or snow), there are two fluxes
to consider, the turbulent and the radiative. The turbulent
fluxes in turn divide into two categories: the latent heat flux
associated with evaporation from leads or sublimation from
the ice (snow) surface and the sensible heat flux due to convective heat transfer over the leads or ice (snow) surface.
The turbulent fluxes are calculated using bulk parameterizations that seek to capture the unresolved role of turbulence in moisture and heat transfer though empirical
relations based on the time-averaged (i.e., bulk) air flow
properties. The radiative fluxes also split into two subcategories: the longwave and the shortwave radiation. The simulated upwelling and downwelling components of the
shortwave radiation are critically dependent on the
description of the surface albedo. The sum of these four
fluxes leads to a surface heat source or sink for the sea
ice, helping to drive either ice melt at the surface or ice
growth at the base (Oberhuber et al., 1993).
A unique aspect of sea-ice modeling, as compared with
land-ice or even ocean modeling, is the treatment of the
mass-conservation equation. While ice is considered
incompressible from a density point of view, similar to
the approximation made in ocean modeling, the threedimensional volume of ice is actually split into two separate
Chapter 16
The Marine Cryosphere
equations: an equation for ice thickness (i.e., the vertical
average over an area) and an equation for the areal fraction,
the area actually covered by sea ice (with the other fraction
denoted as lead area). Taking the combined product of the
ice thickness and areal concentration equation allows one to
reconstitute the original three-dimensional mass (volume)
conservation equation. This mathematical splitting of the
volume (mass) conservation equation into separate equations for thickness and concentration is done as it is natural
to model sea-ice thickness and concentration as separate
variables and also to ease comparisons with observations.
While this is a rather clever split, one complication that
arises is how to apply heat sources and sinks to ice growth
and decay. That is, just as the volume equation has been
split into two equations, the heat source/sink terms need
also to be split, but it is not at all obvious how to partition
the heat source/sink. Specifically, given a unit of heat loss
from the ocean, should that heat be used to grow the ice vertically and extend the thickness or horizontally and extend
the areal coverage? The inverse question can be asked in the
case of a unit of heat gain. Current-generation models have
a relatively ad hoc formulation of how to split this heat flux,
albeit one that generates reasonable model simulations.
Further research is required to better understand and parameterize this heat flux split.
The sea-ice dynamics equations describe how the pack
ice moves under the influence of various forces: ocean
surface tilt, Coriolis effect, wind, currents, and most interestingly, internal ice dynamics, often referred to as ice rheology. The main discussion in theoretical developments of
sea ice centers on an appropriate rheology—or flow law—
that accurately describes how ice floes interact with one
another and how they exchange momentum. Ocean waters
are treated as a Newtonian fluid whereby stress is linearly
related to strain through a fixed viscous coefficient. By contrast, sea ice is thought of as an elastic–plastic fluid
(Rothrock, 1975). At small strain, the stress is described
as elastic and is reversible. At larger strain, the stress is
finite and plastic, meaning the ice fails under tension, compression, or shear stress. An appropriate description, one
that is both physically accurate and numerically efficient,
has been a target of sea-ice dynamicists for decades. Progress has been made, with the current generation of models
simulating ice quite reasonably, starting notably from the
early work of Hibler (1979) in which the observed seaice thickness pattern of the Arctic Ocean, with thick 6-m
ice on the Canadian side and thin 1-m ice on the Russian
side, was reasonably simulated for the first time.
2.3. Ocean Mixed-Layer Interaction
The field of oceanography took a major theoretical step
forward with the discovery of the phenomenon of the
Ekman layer (Ekman, 1905). During his trans-Arctic cruise,
421
Nansen noticed that sea ice did not flow in the direction of
the prevailing wind but drifted somewhat to the right.
Nansen had this observation passed to Ekman who subsequently worked on the mathematical details of how sea
ice and the upper ocean interact. Ekman’s mathematics
described a layer beneath the sea ice that progressively
veered the water current to the right in a spiral fashion, with
deeper water moving slower due to viscous forces. The
underlying dynamics is a simple balance between Coriolis
force and vertical stress. The resulting spiral structure has
been directly observed beneath sea ice (Hunkins, 1966),
as well as at the open ocean surface (Lenn and
Chereskin, 2009) and at the seafloor (Kundu, 1976). The
Ekman layer is of significant dynamical consequence as
it drives a net transport of water to the right (left) of the
surface ocean stress in the northern (southern) hemisphere.
In the context of the marine cryosphere, this has an
important role to play in the upwelling or downwelling of
waters along the margins of the Arctic and Antarctic
(England et al., 1993).
Away from the sea-ice-covered portions of the global
ocean, the surface water layer is mixed down to some depth,
generally in the range 10–100 m. In the resulting surface,
mixed-layer water properties are almost homogeneous in
the vertical (de Boyer Montegut et al., 2004). The key principle underpinning this mixed-layer depth is the balance
between turbulent kinetic energy input and its ability to
overcome the potential energy of the ambient stratification.
The turbulent input comes via surface wind stress, ocean
current shear, and wave breaking, and the opposing
potential energy barrier is simply due to the ambient stratification (Large et al., 1994). In the Arctic Ocean, one of the
most stratified in the world, the top 100–200 m is well stratified with respect to the deeper layers of water, making
exchange between the upper, cold halocline and the deeper,
warmer Atlantic layer quite difficult from an energy point
of view (Aagaard and Carmack, 1989). By contrast, the
Southern Ocean is only weakly stratified and a modest input
of turbulent energy can overcome the weak background
stratification (Martinson, 1990).
Vigorous exchange in the upper ocean, driven by turbulence, rapidly homogenizes scalar properties. In the case of a
sea-ice cover, the exchange between the liquid ocean and
solid ice surface, specifically at the liquid–solid interface,
occurs via molecular transfer processes. The underlying
principle is that turbulent fluid exchange is dampened at a
solid interface as the scales of turbulence decay as one
approaches the solid interface (Kader and Yaglom, 1972).
Extremely close to the interface, the final exchange of scale
properties occurs via molecular (diffusive) processes. At the
ice–ocean interface, both salinity and temperature play a
central role in the melt (growth) rates at the interface. Furthermore, as salt diffusion is approximately 100 times
slower than that of temperature, there is an interesting
422
PART IV
(a)
TS
QTI
Ice
shelf
QTlatent
HI
TB
QTM
Ice–ocean
boundary h
layer
Ocean
mixed
layer
TM
HM
(b)
SI
Ice
shelf
QSI
QSbrine
HI
SB
QSM
Ice–ocean
boundary h
layer
Ocean
mixed
layer
SM
HM
FIGURE 16.4 Schematic representation of the viscous-sublayer model
of ice–ocean interaction showing (a) the heat and (b) the salt balance at
the base of sea ice (or equivalently at an ice-shelf base). The thickness
of ice HI and ocean mixed-layer HM are of the order of meters, while
the ice–ocean boundary layer (i.e., viscous sublayer) h is of the order of
millimeters. Temperatures (salinities) in the ice interior TS (SI ), at the base
TB (SB), and ocean mixed-layer TM (SM) are shown. The heat (salt) fluxes
through the ice QTI (QSI), at the interface QTlatent (QSbrine), and in the ocean QTM
(QSM) are shown. The key control is that the fluxes of heat and salt balance
precisely at the interface. From Holland and Jenkins (1999).
competitive dynamic at the interface. The end result is that
the ice–ocean interface melt (growth) rate is controlled
by the relative rates at which heat and salt can diffuse into
the boundary layer. At the interface, there are three physical
relations to be satisfied simultaneously: the empirical
pressure and salinity-dependent freezing point of sea water,
the balance of heat flux into and out of the interface, and the
balance between the salt flux into the interface and the freshening effect of melt water from the interface. An illustration
of the various scalars and fluxes is presented in Figure 16.4.
This set of constraints gives rise to the “three-equation”
model of ice–ocean interaction (McPhee, 2008).
2.4.
Polynyas
Derived from the Russian word meaning “a large opening in
an otherwise sea-ice covered area,” polynyas occur at
Ocean Circulation and Water Masses
various locations throughout the wintertime sea-ice cover.
Their existence is at first glance surprising, as the surface
energy balance in polar winter is strongly negative such that
the ocean is losing a vast amount of heat, and one would
expect it to be ice covered. Two physical processes that
allow for the occurrence of polynya are, first, the existence
of solid boundaries such as coastline which allow ice to be
blown away from the coast and slowly replenished by new
ice growth (Pease, 1987), and second, vertical convective
processes that can raise warm water from depth (Holland,
2001). In either case, adequate heat is brought into the
mixed layer or ice transported away from the mixed layer,
such that sustained open water areas are possible, even in
wintertime conditions.
From the perspective of the climate modeling and
physical oceanographic communities, recurrent polynyas
are worthy of study (Willmott et al., 2007) because they
are sites where (1) water mass transformation takes place
through the combined effects of cooling and frazil ice formation, (2) large ocean-to-atmosphere heat (several
hundred Watts per square meter) and moisture fluxes occur,
and (3) atmospheric CO2 is sequestered into the ocean by
physical–chemical processes and biological activity.
In coastal polynyas, where the water is shallow, the ocean
quickly cools down to the freezing point at all depths (Lemke,
2001). The heat supplied to the atmosphere then originates
only from the latent heat of fusion due to the continuous production of sea ice. Where offshore winds occur, sea ice is
constantly formed and blown away from the coast, allowing
yet more sea ice to be formed, in effect creating a type of seaice factory. During sea-ice formation, oceanic salt is released
into the water because it is not incorporated into the ice
crystals. This salt increases the local density of the ocean
water, in some cases to the point where it can sink to the deep
ocean. Coastal polynyas therefore represent important
sources of dense deep and bottom water, which ventilate
the abyss and drive the oceanic global conveyor belt.
In open ocean polynyas, the supply of oceanic heat is
less restricted (Lemke, 2001). Therefore, large amounts
of sensible heat from deeper ocean layers are lost to the
atmosphere and little ice is created. The substantial cooling
of the ocean water can lead to deep convection, which
homogenizes the ocean waters to great depths and produces
deep water in the open ocean. Consequently, both types of
polynyas have a large influence on global ocean circulation
and on the earth’s climate.
2.5. Impact on Water Masses, and
Circulation
The mere existence of a sea-ice cover provides a distinctive
surface boundary forcing compared to other regions of
the global ocean. From an oceanographic point of view,
the most important implication of the sea-ice cover is the
Chapter 16
423
The Marine Cryosphere
large input of salt that sea-ice growth inputs to the ocean
surface during winter formation. On the flip side, the
summer melt of the sea ice provides an enormous freshwater input at the surface. Taken together, these two processes result in an effective salt pump that moves salt
from the near surface by convectively driving it down into
the deeper ocean (Broecker and Peng, 1987).
A second impact of the ice cover is a modification to the
input of wind stress (Joffre, 1982). Well known is that the
curl of wind stress is the major input to the Sverdrup balance
governing large-scale ocean circulation. Where there is
extensive sea-ice cover, particularly near coastal regions,
the exchange of momentum between atmosphere and ocean
is fundamentally altered by the existence of the sea-ice
cover. In the case of a lengthy sea-ice margin, the curl of
the wind stress on the ocean can be radically different from
anywhere else in the ocean as there is a sharp line where the
curl abruptly changes (Quadfasel et al., 1987).
3.1.
Observations
Despite the low temperatures of the upper polar oceans, the
presence of sea ice provides a platform that allows one of the
largest and most important components of the earth’s ecosystem to exist. Despite the low light levels, high salinities,
and poor surface gas exchange, microorganisms thrive in this
environment, particularly algae (Lizotte, 2001). As the regulation of carbon dioxide between the ocean and the atmosphere is particularly important to the marine biosphere,
and sea ice plays a key role in modulating gas exchange
and allowing organisms grow, once again the existence of
a sea-ice cover is seen to be important in how the global
ocean interacts with the atmosphere and ultimately the
amount of carbon dioxide stored in the atmosphere, and
hence global air temperatures (Rysgaard et al., 2011).
The global ocean only directly interacts with the ice sheets
where the periphery of the ice terminates in the ocean, either
as an ice shelf, which extends out onto the ocean surface as a
floating tongue, or a tidewater glacier which has an abrupt
transition from land ice to ocean with no floating tongue
(Meier and Post, 1987). It was once thought that the disintegration of an ice shelf, which by definition floats on the
ocean, would have no impact on the flow into the ocean
of the inland ice that fed it (Van der Veen, 1985). That viewpoint has been disproved during the past decade as ice
shelves, such as the Larsen B Ice Shelf, rapidly disintegrated and the flow toward the ocean of the inland ice
behind accelerated manyfold (Scambos et al., 2004).
This knowledge comes from an unprecedented suite of
observational equipment placed in situ on the periphery of
the ice sheets, and airborne and remote sensing instrumentation (Mohr et al., 1998). Remote sensing techniques have
shown a doubling in speed of the Jakobshavn glacier in west
Greenland, following the breakup of the ice tongue in front
of that glacier (Alley et al., 2008). That particular breakup is
believed to have been caused by the sudden arrival of warm
waters into the fjord connecting to the glacier (Holland
et al., 2008). Laser altimetry has captured a drop in surface
elevation over the Greenland outlet glaciers, and their
analogs in the Antarctic, most notably in the Pine Island
Glacier region, located in the Amundsen Sea (Pritchard
et al., 2009). At the same time, correlation techniques
applied to imagery have demonstrated that the velocity of
these outlet glaciers has increased, coincident with the
drop in surface elevations (Joughin et al., 2004). Furthermore, the occurrence of simultaneous change between
neighboring glaciers suggests the ocean to be the main
driver of this change (Shepherd et al., 2004).
3. LAND ICE
3.2.
Land ice, built up through the process of evaporation over
the global ocean and subsequent precipitation over the great
ice sheets of Greenland and Antarctica, has traditionally
been thought to be one of the most passive components
of the climate system and thus of not much consequence
to ocean circulation, particularly on timescales of centuries
or less (Figure 16.5). The pace of interaction was considered to be “glacial.” Since the early 1990s, however,
an unprecedented rapid change in the margins of the ice
sheets has been observed through the eyes of satellites
(Nick et al., 2009), and the cause for change has been
largely deduced to be the global ocean, via the import of
warm waters to the periphery of the ice sheets (Pritchard
et al., 2012). Suddenly, land ice has become a major player
in rapid global sea-level change and water mass modification processes in the polar oceans (Alley et al., 2005).
The physical basis for modeling land ice holds some similarity with sea ice, described in the previous section, but
also has significant differences. A main difference is that
glacial ice is treated as a three-dimensional continuum
(Blatter et al., 2011), with no analog for the description
of concentration (or open space) as is the case for sea ice.
Glacial ice does indeed fracture, but such fractures occupy
a small portion of the total areal coverage and are not yet
widely considered in the overall integrity or strength of
an ice sheet or shelf. There are some efforts to correct this
modeling deficiency by incorporating the concept of
“damage mechanics” into the ice equations (Pralong
et al., 2003).
The thermodynamics of glacial ice is analogous to sea
ice in almost every respect; that is, boundary conditions
apply to the ice surface and base, and molecular diffusion
2.6. Biogeochemical Ramifications
Modeling
424
PART IV
Ocean Circulation and Water Masses
FIGURE 16.5 High-altitude views of the two major ice sheets (a) Greenland and (b) Antarctica. The periphery of the Greenland Ice Sheet interacts with
the surrounding ocean through relatively narrow, deep fjords, some of which have floating tongues and others which terminate abruptly as a vertical face at
the ice–ocean interface. By contrast, the interaction of the Antarctic Ice Sheet with the surrounding ocean is through ice shelves ranging in size from modest
to massive, with the three largest being the Ross, Weddell, and Amery. http://www.climatepedia.org/about-ice-sheets; http://news.discovery.com/earth/
zooms/ocean-warming-melting-antarctica-ice-sheets-120426.html.
of heat is the dominant process in the ice interior (Hutter,
1982). Given the thickness of glacial ice, tending to be
greater than 1 km, it responds to thermal anomalies on
the timescale of hundreds or thousands of years. Sea ice,
on the other hand, is typically of order 1 m thick and
responds to anomalies at the ice–ocean or atmosphere–
ocean boundary rapidly, certainly well less than 1 year.
The dynamics of glacial ice bear little resemblance to
those of sea ice, not only because of the absence of fracture
or void space in the treatment of glacial ice but also because,
unlike sea ice, glacial ice is in general not currently modeled
to fracture. This is quite surprising, given the fact that one of
the most dramatic natural events to occur is the calving, or
fracturing, of large pieces of glacial ice into the coastal
ocean at the margins of the great ice sheets. From the coast,
the icebergs are transported by the mean ocean currents and
winds and subsequently melt into the ocean, providing a
distributed source of freshwater into the ocean. In order
for this process to be adequately captured in global climate
models, significant progress is going to be required in the
physical description of ice calving, and hence of ice rheology. This advance is in turn dependent upon the acquisition of adequate observational data to better understand
the mechanism underlying ice calving.
Chapter 16
425
The Marine Cryosphere
3.3. Ocean Mixed-Layer Interaction
From an oceanographic perspective, the thermodynamic
interaction of the base of an ice shelf with the ocean is
crucial for two reasons: firstly, the melting process modifies
water masses such as Antarctic Bottom Water (AABW)
(Foster and Carmack, 1976), and second, the disintegration
of an ice shelf could eventually contribute to a change in sea
level (Alley et al., 2005). Here we discuss the current
treatment of thermodynamic interaction. The physical
description is almost identical to that of the analogous phenomena at the base of sea ice, as discussed in Section 2. This
is reasonable since the two environments have many
features in common.
One important distinction is that the base of an ice shelf
can be located at considerably higher pressure than that of sea
ice, which is nominally at atmospheric pressure. At depths of
2 km, such as at the grounding zone of the Filchner-Ronne
Ice Shelf in West Antarctica, the pressure increase causes
a lowering of the freezing point to approximately 1.5 C
below the value at the ocean surface (Fujino et al., 1974).
This has the interesting and important consequence that if
water at the surface freezing point (1.9 C), the coldest
water that can be formed by direct interaction with the atmosphere, reaches ice at the deep grounding zones, it finds itself
least 1.5 C warmer than the local freezing point and can
therefore cause high rates of basal melting.
This thermodynamical effect leads to an ocean
dynamical circulation creating the ice pump mechanism
(Figure 16.6). Effectively deep ice is melted because of
the pressure dependence of freezing temperature. The
resulting melt water rises along the base of the ice shelf
Moisture transport
Ice sheet
Glacier
flow
Ice shelf
Refreezing
Sea ice
Iceberg
Ocean
Circumpolar
deep water
Basal
melt
Grounding
zone
Continental shelf
Antarctic
bottom
water
FIGURE 16.6 Schematic representation of the two-dimensional circulation under an ice shelf. Salt rejected by winter sea-ice growth forms
dense, high-salinity water, which sinks and flows beneath the ice shelf.
This causes melt when it comes into contact with deep ice at the grounding
zone. The freshened plume rises under the base of the shelf and can either
refreeze as marine ice or mix with warm, salty Circumpolar Deep Water
(CDW) to form Antarctic Bottom Water (AABW). The circulation is often
referred to as an “ice pump.” From AMISOR (2013).
under the action of gravity because the meltwater is always
more buoyant than the ambient—a consequence of the
equation of state for seawater at low temperatures.
Depending upon the ambient conditions, the rising water
may become supercooled as the pressure lessens and the
freezing point rises, leading to the possibility that frazil
ice may form and be deposited at the base of the ice shelf
(Lewis and Perkins, 1986). This entire cycle in which
ice is melted at depth and then refrozen at shallower iceshelf drafts is thought of as an ice pump circulation—
moving floating ice from deep to shallow elevations at
the ice-shelf base.
In the discussion of sea-ice basal melting in Section 2.3,
the concept of the viscous-sublayer model has been introduced (see again Figure 16.4). This model has also been
adopted for use at the base of an ice shelf (Holland and
Jenkins, 1999), despite not yet being thoroughly validated
for use in such a regime. Other than the greater depth,
another feature distinguishing a sub-ice-shelf base from a
sea-ice base is that an ice-shelf base tends to be inclined,
typically by 0.1% slope, but steeper in some locations, particularly near the grounding zone. Field campaigns are only
now being carried out to make direct measurements of the
turbulent boundary layer at the base of a few Antarctic ice
shelves, with validation and improvement of parameterizations of transport of fluxes through the ice shelf–ocean
boundary layer likely to occur within the next few years.
The first field results suggest that, at least in the instance
of the Pine Island Ice Shelf in West Antarctica, the slope
of the interface and the stratification that results from the
melting along the sloping base form an intense feedback
strengthening the melt rates. This feedback likely contributes to the large-scale channelization of the ice-shelf
base (Stanton et al., 2013).
We have thus far mentioned facets of the ocean mixed
layers beneath a sea-ice cover and also beneath an ice
shelf. Along the front of an actively calving glacier, there
often exists a sea-ice cover with icebergs embedded,
sprinkled through the sea ice. The resulting mélange of
sea ice and iceberg makes for a truly complex ocean
mixed-layer environment, one in which the usual turbulent
kinetic energy input is also influenced by the stirring due
to deep iceberg keels as well as buoyancy input due to
freshwater release as the bergs melt at depth (Burton
et al., 2012). This is an area of ocean mixed-layer
modeling that has not been much addressed to date and
one that is in need of in situ observations upon which a
solid theoretical foundation can be laid.
While sea ice is an obvious form of ice of marine origin,
there is also a form that develops beneath an ice shelf and
can reach impressive thickness. For instance, in the middle
of the Filchner-Ronne Ice Shelf, the thickness of the
bottom-accumulated marine ice layer is estimated to be
almost half the total ice-shelf thickness (Oerter et al.,
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PART IV
1992). This marine ice is formed in the water column as part
of the ice pump mechanism, mentioned above. The
presence of significant thickness of marine ice at the base
of an ice shelf could have significant ramifications for
the strength of the ice shelf as a whole (Holland et al.,
2009) and, if the marine ice survives to the ice front without
being remelted, could play an important role in the manner
in which an ice-shelf front calves.
3.4.
Impacts on Water Masses
The continental shelves of Antarctica, particularly the
western edges of the Ross and Weddell Seas, are wellknown locations for formation of AABW (Foster and
Carmack, 1976). This formation is the result of wintertime
production of sea ice which causes the surface salinities to
reach the highest values found anywhere in the Southern
Ocean. As salinity controls density for waters near the
freezing point, these cold salty surface waters sink under
vertical convection and flow northward down over the
continental shelf break and fill the abyssal global ocean.
Approximately half of the global ocean volume is constituted by this water mass. But as the continental shelves of
Antarctica are foredeepened, some fraction of the dense
water formed over the continental shelves first goes
southward and into the sub-ice-shelf cavity (see again
Figure 16.6). The water that goes beneath the ice shelf generally travels inland all the way back to the grounding zone
where, because of the dependence of the melting point of
ice on pressure, the newly arrived water causes ice-shelf
basal melting. The resulting melt water, known as Ice
Shelf Water (ISW), is buoyant and rises along the ice-shelf
base as mentioned earlier. In some areas, this ISW plays an
important role in AABW production. The production of
much of the AABW over the south western Weddell
Sea continental shelf is regarded as being a result of
ISW mixing with ambient warm, deep water on its way
down the continental slope (Seabrooke et al., 1971). In
the Ross Sea, some significant proportion of AABW is
created by a similar mechanism (Jacobs et al., 1970).
In the eastern Weddell Sea, and in many other parts of
Antarctica, the melt from ice shelves freshens the continental shelf enough to inhibit the production of highsalinity shelf waters, which in turn inhibits the production
of AABW. Turning to sea-ice cover impacts of ISW, it has
been shown in numerical experiments that the mere existence of ISW in the Weddell Sea results in a fresher ocean
surface over the Weddell Sea and consequently there is an
impact on increased sea-ice cover development (Hellmer
et al., 2006).
While dense water formation does occur in the
Greenland Sea, this is truly open-ocean convection and is
Ocean Circulation and Water Masses
not forming along the coast of Greenland where glacierocean interaction might have an influence. Instead, it is
the melting of the Greenland outlet glaciers by warm
waters, from sources distinct from those feeding the convective deep-water formation zones that hold the potential
to impact the convective zones (Aagaard and Carmack,
1989). The formation of deep water in the Greenland Sea
is a major component of the global conveyor belt, and such
formation can in principle be inhibited by the melt water
coming from the Greenland ice sheet (Rahmstorf, 1999).
In the transition between glacial and interglacial periods,
it is thought that the massive freshwater pulse forming from
either the Laurentide or the Greenland Ice Sheets feeds
freshwater that caps the convective zones (Broecker and
Denton, 1990).
3.5. Geochemical Tracers
Understanding the basal melting and water mass modifications beneath an ice shelf has been hampered by the difficulty of accessing the cavity. Setting up sustained
monitoring of the ice-shelf cavity openings has also been
problematic. Aside from traditional CTD (conductivity,
temperature, and depth) observations that can be made
at such locations either through the ice shelf (Nicholls
et al., 1991) or along the ice front (Jacobs et al., 2002),
one can also measure the concentration of various geochemical tracers and, to some degree, deduce the amount
of glacial input to the ocean arising from basal melt
(Smethie and Jacobs, 2005). Additionally, by using measurements of geochemical tracer concentrations to supplement the more common observations of temperature
and salinity, one can better constrain the calculation of
the amount of mixing between various water masses in
the cavity (Schlosser et al., 1990). Another rationale for
using geochemical tracers is that temperature and salinity are “nonconservative” when involved in the melting
of ice in the sense that latent heat and freshwater exchange occur, transforming the two properties. The use
of additional, independent geochemical tracers can remove ambiguity in water mass mixing calculations in such
environments.
From an oceanographic point of view, geochemical
tracers can be either naturally occurring isotopes of, for
instance, oxygen, helium, or neon, or those of anthropogenic origin such as CFCs (chlorofluorocarbons). Over
the past half century, CFCs have been entering the atmosphere and therefore the interior of the ocean through gas
exchange at its surface (Beining and Roether, 1996). For
the purpose of water mass analysis and mixing ratios, if
one knows the amount of CFC in parent water masses,
then one can better estimate the amount of relative mixing
Chapter 16
The Marine Cryosphere
between two or more such water masses based on CFC
measurements in the mixed-water mass. In the instance
of an ice-shelf base, knowing CFC concentration in the
high salinity shelf water (HSSW) that forms in the open
ocean surface adjacent to the ice front and knowing the
CFC concentration in ISW that subsequently forms when
HSSW interacts with the base of an ice shelf, one is able
to deduce the amount of input meltwater. One complicating factor in the case of the ice-shelf regime is that
the presence of a variable sea-ice cover in the formation
region means that equilibration between atmosphere and
ocean is not always established, making the overall
calculations subject to greater uncertainty (Smethie and
Jacobs, 2005).
Other than CFCs, oxygen, helium, and neon have also
been used to study glacial melt into the ocean. Oxygen,
helium, and neon gas are present in glacial ice trapped in
air bubbles or as clathrates (see Section 4). These gases
are released from the ice and dissolve in seawater as a result
of ice-shelf basal melting. The low solubility of helium and
neon in seawater results in concentrations well above solubility equilibrium with the atmosphere, producing a glacial
melt signal that can be readily traced from an ice front,
across the continental shelf and into the abyssal ocean
(Schlosser et al., 1990).
Geochemical tracers aside, another powerful constraint
on calculating water mass mixing and formation in the polar
regions is that relating to the constrained ratio which temperature and salinity change must follow as a water mass
gives up heat to melt ice (Gade, 1979). As ice melts, the
source water doing the melting is cooled and freshened in
a specific ratio depending on physical parameters such as
the latent heat and specific heat capacities of water and
ice. For instance, for the Ross Ice Shelf cavity, an analysis
of the temperature–salinity diagram of the various water
masses in the cavity, coupled with the melting–freezing
(M–F) constraint mentioned above, has led to an improved
quantification of the water mass transformations occurring
there (Figure 16.7).
3.6. Sea-Level Change
On the timescale of the earth’s glacial cycles, the volume of
the global ocean changes substantially, being approximately 100 m or more lower in global sea level during a
glacial period than the sea level of the present interglacial
(Siddall et al., 2003; Chapter 27). This glacial cycle change
in global sea level is thought to be forced by the amount of
solar radiation reaching the earth, the periodicity of which is
known as the Milankovitch cycles, which are dependent
upon the precession, obliquity, and eccentricity of the
427
FIGURE 16.7 A potential temperature–salinity diagram illustrating
melting, freezing, and mixing processes that can occur beneath an ice
shelf. Such diagrams can be useful in identifying the production of Ice
Shelf Water (ISW) and thus inferring the amount of glacial basal melt
into the ocean. By melting and freezing (M–F) at the base of the ice
shelf, the properties of the ocean mixed layer at the base move only
along the solid lines, which have potential temperature to salinity
(Dy/DS) slopes of 2.65. The “þ” symbols represent observed water
masses. The dashed lines indicate possible mixing between observed
varieties of ISW and the high- and low-salinity shelf waters (HSSW
and LSSW). This diagram is for the Ross Ice Shelf cavity. From
Smethie and Jacobs (2005).
earth–sun system (Imbrie et al., 1992). Remarkably, in
these regular cycles of continental ice-sheet growth and
decay, there are periods in which the sea-level changes
abruptly. A remarkable example is the event known as
meltwater-pulse 1A that occurred approximately
14,000 years ago during the last deglaciation (Stanford
et al., 2006). Over a period of a few hundred years, the
global sea level at that time rose by at least 5 m. The implication from this is that future global sea level could also
change as rapidly.
The mechanism by which such rapid past change could
have occurred remains obscure. One theory is that sudden
import of warm ocean waters to the periphery of the ice
shelves, particularly the marine-based portions, leads to a
rapid wasting of the ice that triggers major change over a
short time period. The change in ocean circulation would
have to be driven by changes in atmospheric circulation,
and the mechanism for that change is equally obscure. Continued collection of observational data and development of
coupled global models will likely unravel this mystery in
the coming years.
428
4.
PART IV
MARINE PERMAFROST
In the previous two sections (Section 2 and 3), we have discussed the interaction of ice that floats on the surface of the
ocean with the upper ocean. Ice is also found on the bottom
of the ocean, embedded in the sediment. There are two types
of such ice, one being a pure ice matrix and the other
involving trapped methane.
4.1.
Pure Ice
During glacial periods, large portions of the Siberian Arctic
continental shelves have been covered by ice, and the subsequent flooding of these shelves by ocean waters in interglacial times has resulted in large areas of what was
originally land permafrost being transferred to the marine
permafrost category. In general, marine (or subsea) permafrost is formed either by such inundations or by areas of seafloor with mean annual ocean temperatures below 0 C. A
more technical definition of marine permafrost calls for the
seafloor ice structure to exist for at least a 2-year period
(PAGE21, 2013), a condition easily met along the Siberian
Arctic shelves. The thickness of the marine permafrost layer
has been assessed using direct coring and reaches to 100 m
in some areas. Electromagnetic techniques have also been
employed based on the fact that sediments in an ice-bonded
state have a different electrical resistivity than in the
Ocean Circulation and Water Masses
unbonded state (Corwin, 1983). Seafloor permafrost has
been found only in the Arctic.
4.2. Methane Clathrates
There is a peculiar water-molecule-based structure, formed
under modest pressure (depths greater than 300 m of ocean)
and low temperature (below 0 C), that results in the
trapping of a gas such as methane in a cage surrounded
by water molecules (Max, 2003). The solid is similar to
ice (Figure 16.8a), except for the trapped foreign molecule.
These clathrate hydrates collectively form a component of
the marine permafrost layer. The clathrate can exist at
depths in the marine sediments where the sediment is colder
than the freezing point of the clathrate, itself a function of
pressure (Figure 16.8b). Clathrates are suspected to be
located along the Arctic coastline beneath the marine permafrost that has formed as a result of glacial advances, mentioned above. However, they are not restricted to such
locations. Oceanic deposits seem to be widespread on the
continental shelf and can occur at depth within the sediments or close to the sediment–water interface. They may
cap even larger deposits of gaseous methane (Kvenvolden,
1995). It has been conservatively estimated that the amount
of carbon in the clathrate marine form is more than twice
that found in all fossil fuels on the earth (USGS, 2012).
FIGURE 16.8 (a) A methane clathrate block found in a depth of about 1200 m of water in the upper meter of the sediment. (b) Schematic profile of
temperature through the ocean and uppermost marine sediments. The temperature for clathrate stability, T3(P), increases with pressure (or depth). Experimental data for T3(P) in pure water (dashed line) and seawater (solid line) are shown. The base of the hydrate stability zone (HSZ) is defined by the
intersection of the Geotherm with T3(P). Warming the ocean temperature, and hence the Geotherm, will shallow the HSZ. Panel (a) from http://en.wikipedia.org/wiki/File:Gashydrat_mit_Struktur.jpg; panel (b) from Buffet and Archer (2004).
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The Marine Cryosphere
5. EMERGING CAPABILITIES
In this section, we discuss the emerging capabilities in
observing sea ice and the ocean beneath sea ice and ice
shelves, some of the most difficult parts of the ocean to
study. Improvements and future directions in sea-ice and
land-ice modeling are also given; advances in ocean
modeling are provided elsewhere in Chapter 20.
5.1. Ice-Capable Observations
Starting in the 1970s, the observation of sea ice was revolutionized by satellite remote sensing using passive
microwave instrumentation, leading to unprecedented
sea-ice concentration detail (Gloersen et al., 1992). Subsequent developments have led to sea-ice velocity fields
determined from visual or active radar imagery (Emery
et al., 1997). Now emerging is the possibility of determining
the sea-ice thickness field, the most difficult of the sea-icerelated basic fields to measure by remote sensing (Laxon
et al., 2003). The relevant instruments are the CryoSat radar
altimeter (Wingham et al., 2006) and the ICESat laser
altimeter (Schutz et al., 2005), which are able to give relatively precise observations of sea-ice thickness, and thus
thickness change. This new ability to measure thickness
is a key missing piece needed to more fully observe future
Arctic sea-ice change, as it now allows the possibility,
combined with the sea-ice concentration observations and
sea-ice velocity, to make estimates of sea-ice volume and
mass transport in the coming years.
A number of ice shelves in West Antarctica are undergoing significant mass loss due to what is assumed to be
increasing rates of basal melt (Shepherd et al., 2004). There
has been some efforts to determine basal melt rates using a
combination of remote sensing data sets, basically surface
elevation and surface velocity to infer ice flux divergence,
and thus to evaluate basal melt (Joughin and Padman,
2003). As new satellite instruments come on board over
the next few years, it is expected that observations of melt
rate beneath all ice shelves will improve. However, these
methods do have difficulties when the ice shelf is not in a
steady-state configuration. Satellite altimetric methods
can give an estimate of surface elevation change, which
can then be used for corrections, but these in turn rely on
snow compaction models.
Another important process of ice-shelf mass loss is that
of calving. This is a process that can be well observed from
space, given the number of satellites now in orbit producing
high-resolution imagery (Herried et al., 2011). It will be
possible to form an inventory of icebergs, starting from
the birth to their ultimate demise. Such iceberg tracking
will also provide a data set of freshwater melt input to
the ocean from the icebergs in both polar seas (Gladstone
and Bigg, 2002).
Aside from just taking an inventory of icebergs, there is
still the nagging question of the physical processes that lead
to calving in the first place (Bassis, 2011). To gain insight, it
will be necessary to make measurements of both the strain
and stress fields, in high spatial and high temporal resolution, dictating the need to make such observations in situ,
either on the ice surface or with rapid, repeat airborne
surveys. On the ice, the placement of newly developed
passive seismic instrumentation may turn out to be useful
to map out the stress field before, during, and after a major
calving event (MacAyeal et al., 2009).
5.2.
Ocean-Capable Observations
There are a significant number of emerging technologies to
tackle the notoriously difficult problem of measuring ocean
properties either beneath sea ice or in a sub-ice-shelf cavity.
The subsurface polar ocean is one area where remote sensing
by satellite has not been able to contribute much as the ice
cover provides a surface obstacle to sensing. This implies
that in situ observations are needed throughout the polar
oceans. Perhaps the major breakthrough in this area has been
the successful development and deployment of so-called ITP
(Ice-Tethered-Profilers) (Krishfield et al., 2008; ITP, 2012).
These instruments are profiling CTD devices that travel up
and down a cable passed through access hole drilled through
either sea ice (see Figure 16.9a) or an ice shelf. The data are
retrieved via satellite in almost real time and have enormously increased the observational hydrographic data set
in the Arctic basin (Toole et al., 2011). While the bulk have
been deployed in the Arctic basin to date, their usage is now
spilling over into the Antarctic. The Antarctic, because of its
more seasonal sea-ice cover, is a more challenging environment for ITP deployment.
These ITP instruments are in some ways analogous to
the Argo float program. Argo floats, of course, drift freely
with ocean currents and surface from time to time to
transmit their data (Send et al., 2009), but obviously this
strategy does not work well, or at all, in an ice-covered environment (Klatt et al., 2007). Traditional Argo floats are now
being ice hardened, so they can, if the opportunity arises,
pop up through leads in sea ice (see Figure 16.9b). Another
issue is that Argo floats determine their position by reaching
the ocean surface and using a GPS receiver but, again,
cannot do so in an ice-covered environment. If they are
to operate in a region where they cannot surface often
enough to satisfactorily fix the position of the data they
are measuring, they need to use acoustics to navigate
(Howe and Miller, 2004): arrays of subsurface acoustic
beacons are currently being deployed to provide the necessary acoustic infrastructure.
A shortcoming of both the freely floating ITP and Argo
buoy is that they are not able to be navigated to specific
locations, for example, to perform repeat track surveys.
430
The use of AUV (Autonomous Underwater Vehicles) has
been successful in exploring beneath sea-ice cover and
ice shelves for short periods of time (McPhail et al.,
2009). Such vehicles are now being engineered to have long
range and long duration, including the capability to
PART IV
Ocean Circulation and Water Masses
“hibernate” on the seafloor for extended periods of time.
Although this is a hazardous enterprise, steps can be taken
to minimize the risk (Jalving et al., 2008). For more sustained missions, lasting several months, and possibly up
to a year in the near future, the development of the
FIGURE 16.9 Dominant, emerging observational technologies used to observe the polar ocean beneath the sea-ice and ice-shelf covers. (a) Ice Tethered
Profiler (ITP) anchored to ice, (b) Argo float modified for under-ice deployment,
(Continued)
Chapter 16
The Marine Cryosphere
431
FIGURE 16.9—CONT’D (c) Glider deployed in sea-ice-covered sea, (d) Satellite Relay Data Logger (SRDL) marine-mammal tags attached to a seal,
(e) Distributed Temperature Sensing (DTS) fiber-optic cable prepared for deployment through an Antarctic ice shelf. http://www.polardiscovery.whoi.edu/
expedition1/topic-itp.html; http://www.euro-argo.eu/content/download/21814/314255/file/0830%20Klatt-euroargo-18062010.pdf (Glider photo by
Denise Holland); http://science.nationalgeographic.com/science/photos/antarctic-science-revealed/ (DTS photo by David Holland).
432
PART IV
underwater glider (see Figure 16.9c) perhaps offers the
greatest hope of obtaining hydrographic sections in icecovered oceans (Eichhorn, 2009). Like the Argo buoy for
under-ice missions, the gliders do need navigational assistance from subsurface sound beacons. As an example, such
a glider setup has been recently deployed for Davis Strait in
Baffin Bay (Lee et al., 2004).
The ubiquitous presence of marine mammals in the
polar oceans, such as seals, has led to the development of
the SRDL (Satellite Relay Data Logger). This is a miniature
CTD, coupled with a small Argos transmitter, that is
attached to the back of the neck of a seal (Fedak et al.,
2002). The seal is captured in a net, held for approximately
15 min during which a fast-acting glue, attaches the SRDL
to the seal (see Figure 16.9d) (Teilmann et al., 2000). Such
data sets have the limitation that they only provide data for
approximately 1 year, as the SRDL falls off the seal during
the annual molt. Another issue is that the data are only collected wherever the seal decides to travel. However, as the
behavioral patterns of various seals species is now
beginning to be learned, the SRDL are becoming more
useful for data collection in selected regions of the polar
oceans (Campagna et al., 1999).
A surprising property of a fiber-optical cable is that laser
light shot along the length of the cable can be used to
determine the temperature at approximately every 1 m for
more than a 10-km stretch (Bao et al., 1993). This is possible because the scattered light can be analyzed by a
detector, and based on physical principles of scattering,
along with in-field calibration of the cable, one can obtain
a time series of temperature along the cable at frequent time
and space intervals (Selker et al., 2006). This technique is
known as DTS (Distributed Temperature Sensing) (see
Figure 16.9e for a typical deployment). Such a cable was
recently deployed through the Ross Ice Shelf and a yearlong data set of ice-shelf and sub-ice-shelf temperatures
was collected (Stern et al., 2013).
5.3.
Ice-Capable Modeling
Traditionally, sea-ice models have been developed based on
finite difference discretizations of the governing equations
and solved on either a regular Cartesian or a latitude–
longitude grid (Holland et al., 1993). There are, of course,
much more advanced discretizations and gridding schemes
now developed in other areas of science. Ice researchers
have begun to apply such techniques, for example, finite
element modeling (Timmermann et al., 2009) which allows
for very irregular grids (see Figure 16.10a). Such grids can
be refined to very high resolution in narrow passageways
and channels such as in the Canadian Arctic Archipelago
(Lietaer et al., 2008).
While discretization and gridding dictate the spatial
accuracy of the numerical solutions, time-stepping
Ocean Circulation and Water Masses
efficiency is also a major issue. Global-scale climate
models seek to make each component submodel as efficient
as possible. For the most part, sea-ice models have used
explicit time-integration schemes (CICE, 2012), but more
recent efforts have sought to bring implicit, matrix-based,
methods to bear on the problem. For instance, a matrix-free
Newton–Krylov technique (Lemieux et al., 2012) has been
applied to the sea-ice equations and shows promise for
allowing a significant speed up in integration time.
Numerics aside, aspects of the basic physics of sea ice
still need to be improved. Parameterizations of a host of
subgrid-scale phenomena continue to be tackled by
researchers; for instance, summer melt ponds and their
drainage, brine inclusions, and surface albedo (Holland
et al., 2001). On the dynamics side, there is still uncertainty
regarding the underlying rheology with the earliest rheology focusing on isotropic behavior and more recent
studies looking into the anisotropic behavior of sea ice
(Feltham, 2008). The most common rheology currently in
use is a combination of elastic, viscous, and plastic
behaviors (Hunke and Dukowicz, 1997). One feature still
missing in all these rheologies is that of a tensile strength,
appropriate to the description of landfast ice, as seen along
parts of the coast of the Arctic and Antarctic (König Beatty
and Holland, 2010). Current rheologies have not incorporated this phenomenon, but it is needed as the presence of
landfast ice radically alters the formation of coastal waters,
feeding the deep ocean (Barber and Massom, 2007).
The modeling of land ice, including that of ice shelves,
has lagged significantly behind that of sea ice, but recent
efforts, bolstered by the need for land-ice modeling in IPCC
class models, are beginning to change that situation (CISM,
2012; see Figure 16.10b). The task is significant as it
requires major advances in both the numerical modeling
and the understanding of the physics of land ice and ice
shelves, and also in the coupling of the ice models to ocean
general circulation models. To date, there has not been a
plausible coupling in which the ocean volume and ice
volume swap physical space with one another when either
the ocean melts the ice and invades its physical space or
vice versa (Lipscomb et al., 2009).
6.
CRYOSPHERIC CHANGE
In this final section, we look back on recent changes in the
cryosphere and look ahead to projections for sea ice, land
ice, and marine permafrost in the present century and beyond.
6.1. Observed Sea-Ice Change
The radical reduction in Arctic sea-ice areal extent over the
past few decades has caught the world by surprise. It was
not so long ago, in the early 1990s, when the Arctic sea
ice was thought to be a stable, robust feature of the earth’s
FIGURE 16.10 Examples of emerging technologies in sea-ice and ice-sheet (shelf) modeling. (a) Sea-ice-thickness modeling based on highly variable
grid resolution capable of resolving flow through narrow straits. (b) High-resolution model of the flow of the entire Greenland Ice Sheet. Note the bright red
areas showing the speedup of the ice sheet as it emerges toward the ocean through narrow outlet fjords. https://www.uclouvain.be/en-376557.html; http://
public.lanl.gov/sprice/images/GIS_vels.jpg.
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PART IV
modern climate system, at least on the short timescale of
years to decades. However, satellite remote sensing has
unequivocally shown a nearly 50% decrease in the extent
of the Arctic pack ice during the summer minimum
(Figure 16.11, blue curve). In fact, the past year (2012)
has set an all-time record low (NSIDC, 2012). The cause
of this decline is tied up with the occurrence of global
warming over the past century or so (Stroeve et al.,
2007), as the ice-albedo feedback effect, once it takes hold,
contributes to the accelerated melting. There is of course
natural variability in the Arctic sea-ice cover, as driven,
for example, by the Arctic Oscillation, a pattern of winds that
can either reinforce ice export out of the Artic or hold it
within the basin (NOAA, 2012). Coupled climate models
have been successful in hindcasting Arctic sea-ice decline
in general, but none of the models to date have been able
to accurately hindcast the actual extreme decline, nor the
extreme minima in past years, 2007 and 2012 (Rampal
et al., 2011). This indicates that the coupled models are
not yet fully able to capture extreme events, suggesting more
research is needed either in the individual model components
or in the model couplings to advance this capability.
The story with the Antarctic sea ice is just the opposite,
with a possible small trend upward in extent (Figure 16.11,
red curve), perplexing sea-ice researchers, and global
climate modelers. While it is thought that the ice-albedo
is playing a major role for Arctic sea ice, it is confounding
that no such mechanism appears to be at play in the Antarctic. Of course, a major distinction between the Arctic
and Antarctic is that in the former the sea ice is present
Ocean Circulation and Water Masses
in the summer in appreciable quantity but the Antarctic
sea ice largely disappears, thus possibly reducing its susceptibility to the ice-albedo feedback effect. Additionally,
the Arctic is effectively landlocked by the surrounding
continents, whereas in the Antarctic, the inverse is the
case and this may play a role. There is some evidence
that wind patterns in the Antarctic have controlled the
Antarctic sea ice (Holland and Kwok, 2012). Still, the
apparent lack of an ice-albedo feedback remains a major
question for understanding global climate change and the
cryosphere.
6.2. Sea-Ice Projections
On the short timescale, looking ahead just a few months, an
international initiative called the “Sea Ice Outlook”
(SEARCH, 2012) has joined together a large number of
sea-ice modeling efforts so as to provide a summary of
the expected September Arctic sea-ice minimum. Monthly
reports are released through the summer in an attempt to
provide the scientific community and stakeholders with
the best available information.
On the longer decadal and centennial timescales, sea-ice
hindcasts and projections have been organized under the
IPCC rubric, which has put together simulations based on
ensemble runs of those global climate models that contain
some form of a sea-ice component. The simulations
captured to some degree the climatological annual mean,
seasonal cycle, and temporal trends of sea-ice area during
the modern satellite era, although there was considerable
Arctic and Antarctic Standardized Anomaly and Trend
Nov. 1978 – Dec. 2012
Arctic/Antarctic Standardized Anomaly
(# of st. dev. from 1981 to 2010 average)
4
3
2
1
0
−1
−2
−3
2012
2010
2008
2006
2004
2002
2000
1998
1996
1994
1992
1990
1988
1986
1984
1982
1980
1978
−4
Year
FIGURE 16.11 Arctic (blue curve) and Antarctic (red curve) standardized anomalies and trends of sea-ice cover over the period 1978–2012. Thick lines
indicate 12-month running means, and thin lines indicate monthly anomalies. The Arctic shows a definite trend downward, while the Antarctic shows a
slight increase. http://nsidc.org/cryosphere/sotc/sea_ice.html.
Chapter 16
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The Marine Cryosphere
scatter among the models. Each member of the ensemble
performed differently, and the interannual variability within
each member model is not yet fully understood. An interesting aspect of the simulations is that multimodel ensemble
means show promising estimates close to observations for
the late twentieth century (Zhang and Walsh, 2006). With
regard to sea-ice thickness distributions, considerable deficiencies in sea-ice properties in many model results have
been seen. Most of the models have difficulty in reproducing the spatial sea-ice thickness distribution of the past,
particularly in summer (Gerdes and Köberle, 2007). Taken
collectively, the models indicate a significant reduction in
Arctic sea-ice cover over the coming decades. However,
the large scatter between individual model simulations at
present leads to much uncertainty as to when a seasonally
ice-free Arctic Ocean might be realized (Serreze et al.,
2007). Of real concern is the fact that the models failed
to capture the extreme minimum that occurred in 2007
(prior to the issuance of the IPCC report). An important criterion for future development and improvement of sea-ice
models will be their ability to capture such extreme events.
Projections from IPCC Antarctic sea-ice models show
there is larger uncertainty in model performance in the
Southern Hemisphere (Arzel et al., 2006) with a much
larger disagreement between the model ensemble members
for the Southern Ocean sea-ice simulations. Many of the
models have an annual sea-ice extent cycle that differs
markedly from that observed over the past 30 years
(Turner et al., 2013). In contrast to the satellite data, which
exhibit a slight increase in extent, the mean extent of the
models over the period 1979–2005 shows a decrease
(Turner et al., 2013). Several studies have shown that the
response of the Southern Ocean to an increase in atmospheric greenhouse gas concentrations is delayed compared
with other regions (Goosse and Renssen, 2005), mainly
because of the large thermal inertia of the Southern Ocean.
This highlights the earlier discussion of the bathymetric
contrasts between the Arctic and Southern oceans, the fact
that the former is largely disconnected from the World
Ocean while the latter is not. Aside from an inability to
properly capture the recent ice extent trends, the coupled
models have also not been able to capture the regional
spatial variability of sea-ice cover over the Southern Hemisphere in the past few decades, during which time the Ross
Sea has been gaining cover while the Amundsen has been
losing it (Stammerjohn et al., 2008). Again, this suggests
need for further research to improve sea-ice models, and
their coupling with the ocean and atmosphere.
6.3. Observed Land-Ice Change
The recent breakup of ice shelves in both Greenland (e.g.,
Jakobshavn) and Antarctica (e.g., Larsen B) has also caught
researchers by surprise. The Jakobshavn Ice Shelf has been
undergoing periodic retreat over the past 100 years or so
(Sohn et al., 1998), but with a dramatic thinning and retreat
occurring starting in 1997 (Joughin et al., 2004). No less
dramatic was the breakup of Larsen B, occurring in just a
few weeks in early 2002. From an analysis of the marine
sediments located beneath the former Larsen B Ice Shelf,
it has been demonstrated that the recent collapse has been
unprecedented in the Holocene (Domack et al., 2005).
There is some evidence to suggest that the Greenland
breakup was forced by the arrival of warm waters
(Holland et al., 2008) and that the Antarctic case was driven
by the arrival of warm air over the Antarctic Peninsula
(Marshall et al., 2006). In the case of both Jakobshavn
and the Larsen B ice shelves, the inland glaciers connected
to the shelves have sped up dramatically, with a doubling or
betterment in their speeds (Rignot et al., 2004). The breakup
of these ice shelves has led to a refocus toward both the sensitivity of these ice shelves to climate change and the implications of ice-shelf decay for the stability of grounded ice
(De Angelis and Skvarca, 2003).
A far more disconcerting example of recent glacial
change is the observed thinning of several of the ice shelves
fringing the Amundsen Sea in West Antarctica (Pritchard
et al., 2012). As those ice shelves serve as a front to the
massive marine-based West Antarctic Ice Sheet, any negative mass balance signal in that area could hold serious
consequences for global sea-level rise. Using satellite laser
altimetry combined with modeling of the firn layer, it has
been deduced that ice shelves in this sector of Antarctica
are thinning through increased basal melting (Pritchard
et al., 2012). In fact, the greatest thinning occurs where
warm, deep water can gain access to the continental shelf,
and ultimately the grounding zone of the glaciers, via deep
troughs in the shelves. From ocean modeling studies, it
appears that slowly changing wind stress patterns over
the Amundsen Sea are responsible for the changes in warm
water volume being brought aboard the continental shelf
(Thoma et al., 2008). Understanding what controls the variability of the large-scale wind patterns in the polar regions
and indeed on the global scale is an area of active research
and holds the ultimate key to how warm water can arrive at
the base of an ice shelf (Kerr, 2008). Both Greenland and
Antarctica are showing an overall accelerating mass loss
(Figure 16.12).
6.4.
Land-Ice Projections
The projections of land ice change, and the implicit role that
climate change and warm ocean waters could have on
breaking up of ice shelves and of causing acceleration of
inland ice into the ocean, have been lacking. This shortcoming is implicitly acknowledged by IPCC which failed
to provide an estimate of the potential dynamical contribution of glaciers to sea-level rise over the present century
436
PART IV
FIGURE 16.12 Change in the Antarctic Ice Sheet and Greenland Ice
Sheet during 1992–2010 determined from a reconciliation of measurements acquired by various satellite instrumentation. Also shown is the
equivalent global sea-level contribution (right axis), calculated assuming
that 360 Gt of ice corresponds to 1 mm of sea-level rise. From Shepherd
et al. (2012).
(IPCC, 2007). This apparent shortcoming serves to underscore the difficulty of this task, which the research community has only relatively recently addressed in earnest
(IPCC, 2010). The projections of the contribution to sea
level of the Antarctic Ice Sheet to sea-level rise are
uncertain (Alley et al., 2005) as they range from a few centimeters rise over the next century to a meter or more, the
latter implying devastating consequences for coastal habitats and infrastructure (Nicholls and Cazenave, 2010). Part
of the difficulty in producing credible projections is the lack
of observational evidence upon which to base such model
(Alley et al., 2005). That situation is changing with the
emerging ice and ocean observational technology as discussed in Section 5, but it will take decades to build up
meaningful observational records in these difficult-toaccess regions. Progress in model coupling is likely to
proceed at a faster pace (Lipscomb et al., 2009), but will
have to await the availability of basic observation data,
again highlighting the need for the science community to
have a sustainable ocean and ice observational system in
place around the periphery of the great ice sheets.
6.5.
Marine Permafrost
Concern has been raised that the considerable deposits of
methane, in the form of sub-seafloor clathrates or gas
trapped beneath other permafrost, could be released under
a scenario of global warming leading to a runaway feedback
on global warming. Methane is a powerful greenhouse gas,
having a global warming potential of more than 60 times
that of carbon dioxide. Over the Arctic continental shelf,
methane gas release has indeed been detected (Kort et al.,
2012). The observed distribution of dissolved methane
Ocean Circulation and Water Masses
and possible mechanisms of gas release in connection with
observed dynamics of coastal ocean environments suggest
that areas of the Arctic shelves are important natural sources
of methane to the atmosphere and that such areas tend to be
affected by ongoing global change (Shakhova et al., 2007),
possibly being susceptible to ongoing and future changes in
sea-ice and ocean circulation.
The speculative scenario of a release of vast amounts of
methane or its clathrate form is commonly referred to as the
clathrate gun hypothesis (Kennett et al., 2003), the main
idea being that the seafloor permafrost would itself initially
melt under conditions of a warming global ocean thereby
triggering the release of methane. In its original form, the
hypothesis proposed that the clathrate gun could cause
abrupt runaway warming on a timescale of less than a
human lifetime (Kennett et al., 2003) and might have been
responsible for warming events in and at the end of the last
ice age (Behl, 2000), but more recent research now suggests
this to be unlikely (Sowers, 2006). Past suspected releases
during the Permian–Triassic extinction event and the
Paleocene–Eocene Thermal Maximum have been linked
to such gas release and past climate change (Benton and
Twitchett, 2003).
7.
SUMMARY
In this chapter, we have looked into aspects of the interaction of the World Ocean with the marine cryosphere,
focusing on the Arctic and Antarctic sea-ice covers, the
outlet glaciers of the Greenland and Antarctic ice sheets,
and to a lesser extent, the marine permafrost beneath the
seafloor. We have discussed the fundamental chemical
and physical properties of freezing water and ice that
organize the manner in which ice forms or melts in the
ocean, be it sea ice or land ice, and have seen that heat
exchange at the ice–ocean interface is unique in ocean
dynamics because of the fundamental role played by
molecular exchange. The description of these fine-scale
processes and their encapsulation as parameterizations for
use in global climate models is a maturing subject.
On the large scale, the transport of heat to the polar
ocean is also relatively well understood, but what is not
understood is the manner in which that heat transport
changes through natural variability or anthropogenic
forcing. The global ocean impacts the marine cryosphere
by transporting warm water to the base of sea ice or the base
of an ice shelf, and the ensuing melt impacts back onto the
global ocean through the conveyor belt by both influencing
surface convection and bottom water formation. The ramifications of these large-scale ice–ocean interactions and
potential feedbacks on the climate system are still an area
of active research.
Remote-sensing observational technologies have been
delivering data sets that provide an unprecedented global
Chapter 16
The Marine Cryosphere
view into the marine cryosphere. Perhaps the largest
advance, and surprise, of the past decades has been the witnessing of the enormous change occurring in both sea ice
and land ice. This change has largely reshaped a viewpoint
in which the marine cryosphere was thought to be a relatively fixed feature of the climate system on the timescale
of decades or a century. We have now begun to see that
viewpoint unravel and be replaced by one where the new
norm is Arctic sea ice being halved in summer extent and
a number of outlet glaciers doubling their seaward velocity.
Such change shows the sensitivity of the marine cryosphere
to climate change, whether natural or anthropogenic, and
the pressing need for further sub-ice, polar ocean observational data sets and numerical modeling to provide a
credible and robust ability to project the future of the marine
cryosphere.
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