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This chapter should be cited as: Vaughan, D.G., J.C. Comiso, I. Allison, J. Carrasco, G. Kaser, R. Kwok, P. Mote, T. Murray, F. Paul, J. Ren, E. Rignot, O. Solomina, K. Steffen and T. Zhang, 2013: Observations: Cryosphere. In: Climate Change 2013: The Physical Sci- ence Basis. Contribution of Working Group I to the Fifth Assessment Report of the Intergovernmental Panel on Climate Change [Stocker, T.F., D. Qin, G.-K. Plattner, M. Tignor, S.K. Allen, J. Boschung, A. Nauels, Y. Xia, V. Bex and P.M. Midgley (eds.)]. Cambridge University Press, Cambridge, United Kingdom and New York, NY, USA.
Coordinating Lead Authors: David G. Vaughan (UK), Josefino C. Comiso (USA) Lead Authors: Ian Allison (Australia), Jorge Carrasco (Chile), Georg Kaser (Austria/Italy), Ronald Kwok (USA), Philip Mote (USA), Tavi Murray (UK), Frank Paul (Switzerland/Germany), Jiawen Ren (China), Eric Rignot (USA), Olga Solomina (Russian Federation), Konrad Steffen (USA/Switzerland), Tingjun Zhang (USA/China) Contributing Authors: Anthony A. Arendt (USA), David B. Bahr (USA), Michiel van den Broeke (Netherlands), Ross Brown (Canada), J. Graham Cogley (Canada), Alex S. Gardner (USA), Sebastian Gerland (Norway), Stephan Gruber (Switzerland), Christian Haas (Canada), Jon Ove Hagen (Norway), Regine Hock (USA), David Holland (USA), Matthias Huss (Switzerland), Thorsten Markus (USA), Ben Marzeion (Austria), Rob Massom (Australia), Geir Moholdt (USA), Pier Paul Overduin (Germany), Antony Payne (UK), W. Tad Pfeffer (USA), Terry Prowse (Canada), Valentina Radić (Canada), David Robinson (USA), Martin Sharp (Canada), Nikolay Shiklomanov (USA), Sharon Smith (Canada), Sharon Stammerjohn (USA), Isabella Velicogna (USA), Peter Wadhams (UK), Anthony Worby (Australia), Lin Zhao (China) Review Editors: Jonathan Bamber (UK), Philippe Huybrechts (Belgium), Peter Lemke (Germany)
Observations: Cryosphere
Table of Contents
Executive Summary ..................................................................... 319 4.1 Introduction ...................................................................... 321 4.2 Sea Ice ................................................................................ 323 4.2.1 Background ............................................................... 323 4.2.2 Arctic Sea Ice ............................................................ 323 4.2.3 Antarctic Sea Ice ....................................................... 330 4.3 Glaciers ............................................................................... 335 4.3.1 Current Area and Volume of Glaciers ........................ 335 4.3.2 Methods to Measure Changes in Glacier Length, Area and Volume/Mass ............................................. 335 4.3.3 Observed Changes in Glacier Length, Area and Mass .................................................................. 338 4.4 Ice Sheets .......................................................................... 344 4.4.1 Background ............................................................... 344 4.4.2 Changes in Mass of Ice Sheets .................................. 344 4.4.3 Total Ice Loss from Both Ice Sheets ........................... 353 4.4.4 Causes of Changes in Ice Sheets ............................... 353 4.4.5 Rapid Ice Sheet Changes........................................... 355 4.5 Seasonal Snow ................................................................. 358 4.5.1 Background ............................................................... 358 4.5.2 Hemispheric View...................................................... 358 4.5.3 Trends from In Situ Measurements............................ 359 4.5.4 Changes in Snow Albedo .......................................... 359 Box 4.1: Interactions of Snow within the Cryosphere .................................................................... 360 4.6 Lake and River Ice ........................................................... 361 4.7 Frozen Ground .................................................................. 362 4.7.1 Background ............................................................... 362 4.7.2 Changes in Permafrost .............................................. 362 4.7.3 Subsea Permafrost .................................................... 364 4.7.4 Changes in Seasonally Frozen Ground ...................... 364
4.8 Synthesis ............................................................................ 367 References .................................................................................. 369 Appendix 4.A: Details of Available and Selected Ice Sheet Mass Balance Estimates from 1992 to 2012 ........... 380 Frequently Asked Questions FAQ 4.1 How Is Sea Ice Changing in the Arctic and Antarctic? ........................................................ 333 FAQ 4.2 Are Glaciers in Mountain Regions Disappearing? ..............................................................x Supplementary Material Supplementary Material is available in online versions of the report.
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Ice Sheets The Greenland ice sheet has lost ice during the last two decades (very high confidence). Combinations of satellite and airborne remote sensing together with field data indicate with high confidence that the ice loss has occurred in several sectors and that large rates of mass loss have spread to wider regions than reported in AR4. {4.4.2, 4.4.3, Figures 4.13, 4.15, 4.17} The rate of ice loss from the Greenland ice sheet has accelerated since 1992. The average rate has very likely increased from 34 [–6 to 74] Gt yr–1^ over the period 1992–2001 (sea level equivalent, 0.09 [–0.02 to 0.20] mm yr–1), to 215 [157 to 274] Gt yr–1^ over the period 2002–2011 (0.59 [0.43 to 0.76] mm yr–1). {4.4.3, Figures 4.15, 4.17} Ice loss from Greenland is partitioned in approximately similar amounts between surface melt and outlet glacier discharge (medium confidence), and both components have increased (high confidence). The area subject to summer melt has increased over the last two decades (high confidence). {4.4.2} The Antarctic ice sheet has been losing ice during the last two decades (high confidence). There is very high confidence that these losses are mainly from the northern Antarctic Peninsula and the Amundsen Sea sector of West Antarctica, and high confidence that they result from the acceleration of outlet glaciers. {4.4.2, 4.4.3, Figures 4.14, 4.16, 4.17} The average rate of ice loss from Antarctica likely increased from 30 [–37 to 97] Gt yr–1^ (sea level equivalent, 0.08 [–0.10 to 0.27] mm yr–1) over the period 1992–2001, to 147 [72 to 221] Gt yr–1^ over the period 2002–2011 (0.40 [0.20 to 0.61] mm yr–1). {4.4.3, Figures 4.16, 4.17} In parts of Antarctica, floating ice shelves are undergoing substantial changes (high confidence). There is medium confidence that ice shelves are thinning in the Amundsen Sea region of West Antarctica, and medium confidence that this is due to high ocean heat flux. There is high confidence that ice shelves round the Antarctic Peninsula continue a long-term trend of retreat and partial collapse that began decades ago. {4.4.2, 4.4.5} Snow Cover Snow cover extent has decreased in the Northern Hemisphere, especially in spring (very high confidence). Satellite records indi- cate that over the period 1967–2012, annual mean snow cover extent decreased with statistical significance; the largest change, –53% [very likely, –40% to –66%], occurred in June. No months had statistically significant increases. Over the longer period, 1922–2012, data are available only for March and April, but these show a 7% [very likely, 4.5% to 9.5%] decline and a strong negative [–0.76] correlation with March–April 40°N to 60°N land temperature. {4.5.2, 4.5.3}
Station observations of snow, nearly all of which are in the Northern Hemisphere, generally indicate decreases in spring, especially at warmer locations (medium confidence). Results depend on station elevation, period of record, and variable measured (e.g., snow depth or duration of snow season), but in almost every study surveyed, a majority of stations showed decreasing trends, and stations at lower elevation or higher average temperature were the most liable to show decreases. In the Southern Hemisphere, evidence is too limited to conclude whether changes have occurred. {4.5.2, 4.5.3, Figures 4.19, 4.20, 4.21} Freshwater Ice The limited evidence available for freshwater (lake and river) ice indicates that ice duration is decreasing and average seasonal ice cover shrinking (low confidence). For 75 Northern Hemisphere lakes, for which trends were available for 150-, 100- and 30-year peri- ods ending in 2005, the most rapid changes were in the most recent period (medium confidence), with freeze-up occurring later (1.6 days per decade) and breakup earlier (1.9 days per decade). In the North American Great Lakes, the average duration of ice cover declined 71% over the period 1973–2010. {4.6} Frozen Ground Permafrost temperatures have increased in most regions since the early 1980s (high confidence) although the rate of increase has varied regionally. The temperature increase for colder perma- frost was generally greater than for warmer permafrost (high confi- dence). {4.7.2, Table 4.8, Figure 4.24} Significant permafrost degradation has occurred in the Russian European North (medium confidence). There is medium confidence that, in this area, over the period 1975–2005, warm permafrost up to 15 m thick completely thawed, the southern limit of discontinuous per- mafrost moved north by up to 80 km and the boundary of continuous permafrost moved north by up to 50 km. {4.7.2} In situ measurements and satellite data show that surface sub- sidence associated with degradation of ice-rich permafrost occurred at many locations over the past two to three decades (medium confidence). {4.7.4} In many regions, the depth of seasonally frozen ground has changed in recent decades (high confidence). In many areas since the 1990s, active layer thicknesses increased by a few centimetres to tens of centimetres (medium confidence). In other areas, especially in northern North America, there were large interannual variations but few significant trends (high confidence). The thickness of the season- ally frozen ground in some non-permafrost parts of the Eurasian conti- nent likely decreased, in places by more than 30 cm from 1930 to 2000 (high confidence) {4.7.4}
Observations: Cryosphere Chapter 4
4.1 Introduction
The cryosphere is the collective term for the components of the Earth system that contain a substantial fraction of water in the frozen state (Table 4.1). The cryosphere comprises several components: snow, river and lake ice; sea ice; ice sheets, ice shelves, glaciers and ice caps; and frozen ground which exist, both on land and beneath the oceans (see Glossary and Figure 4.1). The lifespan of each component is very differ- ent. River and lake ice, for example, are transient features that general- ly do not survive from winter to summer; sea ice advances and retreats with the seasons but especially in the Arctic can survive to become multi-year ice lasting several years. The East Antarctic ice sheet, on the other hand, is believed to have become relatively stable around 14 million years ago (Barrett, 2013). Nevertheless, all components of the cryosphere are inherently sensitive to changes in air temperature and precipitation, and hence to a changing climate (see Chapter 2).
Changes in the longer-lived components of the cryosphere (e.g., glaciers) are the result of an integrated response to climate, and the cryosphere is often referred to as a ‘natural thermometer’. But as our understanding of the complexity of this response has grown, it is increasingly clear that elements of the cryosphere should rather be considered as a ‘natural
Ice on Land Percent of Global Land Surfacea^ Sea Level Equivalentb^ (metres) Antarctic ice sheetc^ 8.3 58. Greenland ice sheetd^ 1.2 7. Glacierse^ 0.5 0. Terrestrial permafrostf^ 9–12 0.02–0.10g Seasonally frozen groundh^33 Not applicable Seasonal snow cover (seasonally variable)i 1.3–30.6 0.001–0. Northern Hemisphere freshwater (lake and river) icej^ 1.1 Not applicable Totalk^ 52.0–55.0% ~66. Ice in the Ocean Percent of Global Ocean Areaa^ Volumel^ (10^3 km^3 ) Antarctic ice shelves 0.45m^ ~ Antarctic sea ice, austral summer (spring)n^ 0.8 (5.2) 3.4 (11.1) Arctic sea ice, boreal autumn (winter/spring)n^ 1.7 (3.9) 13.0 (16.5) Sub-sea permafrosto^ ~0.8 Not available Totalp^ 5.3–7.
climate-meter’, responsive not only to temperature but also to other climate variables (e.g., precipitation). However, it remains the case that the conspicuous and widespread nature of changes in the cryosphere (in particular, sea ice, glaciers and ice sheets) means these changes are frequently used emblems of the impact of changing climate. It is thus imperative that we understand the context of current change within the framework of past changes and natural variability. The cryosphere is, however, not simply a passive indicator of climate change; changes in each component of the cryosphere have a signifi- cant and lasting impact on physical, biological and social systems. Ice sheets and glaciers exert a major control on global sea level (see Chap- ters 5 and 13), ice loss from these systems may affect global ocean circulation and marine ecosystems, and the loss of glaciers near popu- lated areas as well as changing seasonal snow cover may have direct impacts on water resources and tourism (see WGII Chapters 3 and 24). Similarly, reduced sea ice extent has altered, and in the future may continue to alter, ocean circulation, ocean productivity and regional climate and will have direct impacts on shipping and mineral and oil exploration (see WGII, Chapter 28). Furthermore, decline in snow cover and sea ice will tend to amplify regional warming through snow and ice-albedo feedback effects (see Glossary and Chapter 9). In addition,
Table 4.1 | Representative statistics for cryospheric components indicating their general significance.
Notes: a b Assuming a global land area of 147.6 Mkm^2 and ocean area of 362.5 Mkm^2. c See Glossary. Assuming an ice density of 917 kg m–3, a seawater density of 1028 kg m–3, with seawater replacing ice currently below sea level. d Area of grounded ice sheet not including ice shelves is 12.295 Mkm^2 (Fretwell et al., 2013). e Area of ice sheet and peripheral glaciers is 1.801 Mkm^2 (Kargel et al., 2012). SLE (Bamber et al., 2013). f Calculated from glacier outlines (Arendt et al., 2012), includes glaciers around Greenland and Antarctica. For sources of SLE see Table 4.2. g Area of permafrost excluding permafrost beneath the ice sheets is 13.2 to 18.0 Mkm^2 (Gruber, 2012). h Value indicates the full range of estimated excess water content of Northern Hemisphere permafrost (Zhang et al., 1999). i Long-term average maximum of seasonally frozen ground is 48.1 Mkm^2 (Zhang et al., 2003); excludes Southern Hemisphere. j Northern Hemisphere only (Lemke et al., 2007). k Areas and volume of freshwater (lake and river ice) were derived from modelled estimates of maximum seasonal extent (Brooks et al., 2012). l To allow for areas of permafrost and seasonally frozen ground that are also covered by seasonal snow, total area excludes seasonal snow cover. m Antarctic austral autumn (spring) (Kurtz and Markus, 2012); and Arctic boreal autumn (winter) (Kwok et al., 2009). For the Arctic, volume includes only sea ice in the Arctic Basin. n Area is 1.617 Mkm^2 (Griggs and Bamber, 2011). o Maximum and minimum areas taken from this assessment, Sections 4.2.2 and 4.2.3. Few estimates of the area of sub-sea permafrost exist in the literature. The estimate shown, 2.8 Mkm (2012).^2 , has significant uncertainty attached and was assembled from other publications by Gruber p (^) Summer and winter totals assessed separately.
Observations: Cryosphere Chapter 4
changes in frozen ground (in particular, ice-rich permafrost) will damage some vulnerable Arctic infrastructure (see WGII, Chapter 28), and could substantially alter the carbon budget through the release of methane (see Chapter 6).
Since AR4, substantial progress has been made in most types of cry- ospheric observations. Satellite technologies now permit estimates of regional and temporal changes in the volume and mass of the ice sheets. The longer time series now available enable more accu- rate assessments of trends and anomalies in sea ice cover and rapid identification of unusual events such as the dramatic decline of Arctic summer sea ice extent in 2007 and 2012. Similarly, Arctic sea ice thick- ness can now be estimated using satellite altimetry, allowing pan-Arc- tic measurements of changes in volume and mass. A new global glacier inventory includes nearly all glaciers (Arendt et al., 2012) (42% in AR4) and allows for much better estimates of the total ice volume and its past and future changes. Remote sensing measurements of regional glacier volume change are also now available widely and modelling of glacier mass change has improved considerably. Finally, fluctuations in the cryosphere in the distant and recent past have been mapped with increasing certainty, demonstrating the potential for rapid ice loss, compared to slow recovery, particularly when related to sea level rise.
This chapter describes the current state of the cryosphere and its indi- vidual components, with a focus on recent improvements in under- standing of the observed variability, changes and trends. Projections of future cryospheric changes (e.g., Chapter 13) and potential drivers (Chapter 10) are discussed elsewhere. Earlier IPCC reports used cry- ospheric terms that have specific scientific meanings (see Cogley et al., 2011), but have rather different meanings in everyday language. To avoid confusion, this chapter uses the term ‘glaciers’ for what was previously termed ‘glaciers and ice caps’ (e.g., Lemke et al., 2007). For the two largest ice masses of continental size, those covering Green- land and Antarctica, we use the term ‘ice sheets’. For simplicity, we use units such as gigatonnes (Gt, 10^9 tonnes, or 10^12 kg). One gigatonne is approximately equal to one cubic kilometre of freshwater (1.1 km^3 of ice), and 362.5 Gt of ice removed from the land and immersed in the oceans will cause roughly 1 mm of global sea level rise (Cogley, 2012).
4.2 Sea Ice
4.2.1 Background
Sea ice (see Glossary) is an important component of the climate system. A sea ice cover on the ocean changes the surface albedo, insu- lates the ocean from heat loss, and provides a barrier to the exchange of momentum and gases such as water vapour and CO 2 between the ocean and atmosphere. Salt ejected by growing sea ice alters the den- sity structure and modifies the circulation of the ocean. Regional cli- mate changes affect the sea ice characteristics and these changes can feed back on the climate system, both regionally and globally. Sea ice is also a major component of polar ecosystems; plants and animals at all trophic levels find a habitat in, or are associated with, sea ice.
Most sea ice exists as pack ice, and wind and ocean currents drive the drift of individual pieces of ice (called floes). Divergence and shear in
sea ice motion create areas of open water where, during colder months, new ice can quickly form and grow. On the other hand, convergent ice motion causes the ice cover to thicken by deformation. Two relatively thin floes colliding with each other can ‘raft’, stacking one on top of the other and thickening the ice. When thicker floes collide, thick ridges may be built from broken pieces, with a height above the surface (ridge sail) of 2 m or more, and a much greater thickness (~10 m) and width below the ocean surface (ridge keel). Sea ice thickness also increases by basal freezing during winter months. But the thicker the ice becomes the more it insulates heat loss from the ocean to the atmosphere and the slower the basal growth is. There is an equilibrium thickness for basal ice growth that is dependent on the surface energy balance and heat from the deep ocean below. Snow cover lying on the surface of sea ice provides additional insulation, and also alters the surface albedo and aerodynamic roughness. But also, and particularly in the Antarctic, a heavy snow load on thin sea ice can depress the ice surface and allow seawater to flood the snow. This saturated snow layer freezes quickly to form ‘snow ice’ (see FAQ 4.1). Because sea ice is formed from seawater it contains sea salt, mostly in small pockets of concentrated brine. The total salt content in newly formed sea ice is only 25 to 50% of that in the parent seawater, and the residual salt rejected as the sea ice forms alters ocean water density and stability. The salinity of the ice decreases as it ages, and particu- larly during the Arctic summer when melt water (including from melt ponds that form on the surface) drains through and flushes the ice. The salinity and porosity of sea ice affect its mechanical strength, its thermal properties and its electrical properties – the latter being very important for remote sensing. Geographical constraints play a dominant but not an exclusive role in determining the quite different characteristics of sea ice in the Arctic and the Antarctic (see FAQ 4.1). This is one of the reasons why changes in sea ice extent and thickness are very different in the north and the south. We also have much more information on Arctic sea ice thickness than we do on Antarctic sea ice thickness, and so discuss Arctic and Antarctic separately in this assessment. 4.2.2 Arctic Sea Ice Regional sea ice observations, which span more than a century, have revealed significant interannual changes in sea ice coverage (Walsh and Chapman, 2001). Since the advent of satellite multichannel pas- sive microwave imaging systems in 1979, which now provide more than 34 years of continuous coverage, it has been possible to monitor the entire extent of sea ice with a temporal resolution of less than a day. A number of procedures have been used to convert the observed microwave brightness temperature into sea ice concentration— the fractional area of the ocean covered by ice—and thence to derive sea ice extent and area (Markus and Cavalieri, 2000; Comiso and Nishio, 2008). Sea ice extent is defined as the sum of ice covered areas with concentrations of at least 15%, while ice area is the product of the ice concentration and area of each data element within the ice extent. A brief description of the different techniques for deriving sea ice con- centration is provided in the Supplementary Material. The trends in the sea ice concentration, ice extent and ice area, as inferred from data
Chapter 4 Observations: Cryosphere
derived from the different techniques, are generally compatible. A com- parison of derived ice extents from different sources is presented in the next section and in the Supplementary Material. Results presented in this assessment are based primarily on a single technique (Comiso and Nishio, 2008) but the use of data from other techniques would provide generally the same conclusions. Arctic sea ice cover varies seasonally, with average ice extent varying between about 6 × 10^6 km^2 in the summer and about 15 × 10^6 km^2 in the winter (Comiso and Nishio, 2008; Cavalieri and Parkinson, 2012; Meier et al., 2012). The summer ice cover is confined to mainly the Arctic Ocean basin and the Canadian Arctic Archipelago, while winter sea ice reaches as far south as 44°N, into the peripheral seas. At the end of summer, the Arctic sea ice cover consists primarily of the pre- viously thick, old and ridged ice types that survived the melt period. Interannual variability is largely determined by the extent of the ice cover in the peripheral seas in winter and by the ice cover that survives the summer melt in the Arctic Basin. 4.2.2.1 Total Arctic Sea Ice Extent and Concentration Figure 4.2 (derived from passive microwave data) shows both the sea- sonality of the Arctic sea ice cover and the large decadal changes that have occurred over the last 34 years. Typically, Arctic sea ice reaches its maximum seasonal extent in February or March whereas the min- imum occurs in September at the end of summer melt. Changes in decadal averages in Arctic ice extent are more pronounced in summer than in winter. The change in winter extent between 1979–1988 and 1989–1998 was negligible. Between 1989–1998 and 1999–2008, there was a decrease in winter extent of around 0.6 × 10^6 km^2. This can be contrasted to a decrease in ice extent at the end of the summer (September) of 0.5 × 10^6 km^2 between 1979–1988 and 1989–1998, followed by a further decrease of 1.2 × 10^6 km^2 between 1989– and 1999–2008. Figure 4.2 also shows that the change in extent from 1979–1988 to 1989–1998 was statistically significant mainly in spring and summer while the change from 1989–1998 to 1999–2008 was statistically significant during winter and summer. The largest inter- annual changes occur during the end of summer when only the thick components of the winter ice cover survive the summer melt (Comiso et al., 2008; Comiso, 2012). For comparison, the average extents during the 2009–2012 period are also presented: the extent during this period was considerably less than in earlier periods in all seasons, except spring. The summer min- imum extent was at a record low in 2012 following an earlier record set in 2007 (Stroeve et al., 2007; Comiso et al., 2008). The minimum ice extent in 2012 was 3.44 × 10^6 km^2 while the low in 2007 was 4. × 10^6 km^2. For comparison, the record high value was 7.86 × 10^6 km^2 in 1980. The low extent in 2012 (which is 18.5% lower than in 2007) was probably caused in part by an unusually strong storm in the Cen- tral Arctic Basin on 4 to 8 August 2012 (Parkinson and Comiso, 2013). The extents for 2007 and 2012 were almost the same from June until the storm period in 2012, after which the extent in 2012 started to trend considerably lower than in 2007. The error bars, which represent 1 standard deviation (1σ) of samples used to estimate each data point, are smallest in the first decade and get larger with subsequent decades indicating much higher interannual variability in recent years. The error
bars are also comparable in summer and winter during the first decade but become progressively larger for summer compared to winter in subsequent decades. These results indicate that the largest interannual variability has occurred in the summer and in the recent decade. Although relatively short as a climate record, the 34-year satellite record is long enough to allow determination of significant and con- sistent trends of the time series of monthly anomalies (i.e., difference between the monthly and the averages over the 34-year record) of ice extent, area and concentration. The trends in ice concentration for the winter, spring, summer and autumn for the period November 1978 to December 2012 are shown in Figure 4.2 (b, c, d and e). The seasonal trends for different regions, except the Bering Sea, are negative. Ice cover changes are relatively large in the eastern Arctic Basin and most peripheral seas in winter and spring, while changes are pronounced almost everywhere in the Arctic Basin, except at greater than 82°N, in summer and autumn. In connection with a comprehensive observation- al research program during the International Polar Year 2007–2008, regional studies primarily on the Canadian side of the Arctic revealed very similar patterns of spatial and interannual variability of the sea ice cover (Derksen et al., 2012). From the monthly anomaly data, the trend in sea ice extent in the Northern Hemisphere (NH) for the period from November 1978 to December 2012 is –3.8 ± 0.3% per decade (very likely) (see FAQ 4.1). The error quoted is calculated from the standard deviation of the slope of the regression line. The baseline for the monthly anomalies is the average of all data for each month from November 1978 to December
- The trends for different regions vary greatly, ranging from +7.3% per decade in the Bering Sea to –13.8% per decade in the Gulf of St. Lawrence. This large spatial variability is associated with the complex- ity of the atmospheric and oceanic circulation system as manifested in the Arctic Oscillation (Thompson and Wallace, 1998). The trends also differ with season (Comiso and Nishio, 2008; Comiso et al., 2011). For the entire NH, the trends in ice extent are –2.3 ± 0.5%, –1.8 ± 0.5%, –6.1 ± 0.8% and –7.0 ± 1.5% per decade (very likely) in winter, spring, summer and autumn, respectively. The corresponding trends in ice area are –2.8 ± 0.5%, –2.2 ± 0.5%, –7.2 ± 1.0%, and –7.8 ± 1.3% per decade (very likely). Similar results were obtained by (Cavalieri and Parkinson, 2012) but cannot be compared directly since their data are for the period from 1979 to 2010 (see Supplementary Material). The trends for ice extent and ice area are comparable except in the summer and autumn, when the trend in ice area is significantly more than that in ice extent. This is due in part to increasing open water areas within the pack that may be caused by more frequent storms and more diver- gence in the summer (Simmonds et al., 2008). The trends are larger in the summer and autumn mainly because of the rapid decline in the multi-year ice cover (Comiso, 2012), as discussed in Section 4.2.2.3. The trends in km^2 yr–1^ were estimated as in Comiso and Nishio (2008) and Comiso (2012) but the percentage trends presented in this chapter were calculated differently. Here the percentage is calculated as a dif- ference from the first data point on the trend line whereas the earlier estimations used the difference from the mean value. The new percent- age trends are only slightly different from the previous ones and the conclusions about changes are the same.
Chapter 4 Observations: Cryosphere
1979–2012 are shown in Figure 4.4. Perennial ice is that which survives the summer, and the ice extent at summer minimum has been used as a measure of its coverage (Comiso, 2002). Multi-year ice (as defined by World Meteorological Organization) is ice that has survived at least two summers. Generally, multi-year ice is less saline and has a distinct microwave signature that differs from the seasonal ice, and thus can be discriminated and monitored with satellite microwave radiometers (Johannessen et al., 1999; Zwally and Gloersen, 2008; Comiso, 2012). Figure 4.4 shows similar interannual variability and large trends for both perennial and multi-year ice for the period 1979 to 2012. The extent of the perennial ice cover, which was about 7.9 × 10^6 km^2 in 1980, decreased to as low as 3.5 × 10^6 km^2 in 2012. Similarly, the multi-year ice extent decreased from about 6.2 × 10^6 km^2 in 1981 to about 2.5 × 10^6 km^2 in 2012. The trends in perennial ice extent and
Figure 4.3 | Ice extent in the Arctic from 1870 to 2011. (a) Annual ice extent and (b) seasonal ice extent using averages of mid-month values derived from in situ and other sources including observations from the Danish meteorological stations from 1870 to 1978 (updated from, Walsh and Chapman, 2001). Ice extent from a joint Hadley and National Oceanic and Atmospheric Administration (NOAA) project (called HADISST1_ Ice) from 1900 to 2011 is also shown. The yearly and seasonal averages for the period from 1979 to 2011 are shown as derived from Scanning Multichannel Microwave Radi- ometer (SMMR) and Special Sensor Microwave/Imager (SSM/I) passive microwave data using the Bootstrap Algorithm (SBA) and National Aeronautics and Space Administra- tion (NASA) Team Algorithm, Version 1 (NT1), using procedures described in Comiso and Nishio (2008), and Cavalieri et al. (1984), respectively; and from Advanced Microwave Scanning Radiometer, Version 2 (AMSR2) using algorithms called AMSR Bootstrap Algo- rithm (ABA) and NASA Team Algorithm, Version 2 (NT2), described in Comiso and Nishio (2008) and Markus and Cavalieri (2000). In (b), data from the different seasons are shown in different colours to illustrate variation between seasons, with SBA data from the procedure in Comiso and Nishio (2008) shown in black.
Figure 4.4 | Annual perennial (blue) and multi-year (green) sea ice extent (a) and sea ice area (b) in the Central Arctic from 1979 to 2012 as derived from satellite passive microwave data (updated from Comiso, 2012). Perennial ice values are derived from summer minimum ice extent, while the multi-year ice values are averages of those from December, January and February. The gold lines (after 2002) are from AMSR-E data. Uncertainties in the observations (very likely range) are indicated by representative error bars, and uncertainties in the trends are given (very likely range).
ice area were strongly negative at –11.5 ± 2.1 and –12.5 ± 2.1% per decade (very likely) respectively. These values indicate an increased rate of decline from the –6.4% and –8.5% per decade, respectively, reported for the 1979 to 2000 period by Comiso (2002). The trends in multi-year ice extent and area are even more negative, at –13.5 ± 2. and –14.7 ± 3.0% per decade (very likely), respectively, as updated for the period 1979 to 2012 (Comiso, 2012). The more negative trend in ice area than in ice extent indicates that the average ice concentration of multi-year ice in the Central Arctic has also been declining. The rate of decline in the extent and area of multi-year ice cover is consistent with the observed decline of old ice types from the analysis of ice drift and ice age by Maslanik et al. (2007), confirming that older and thicker ice types in the Arctic have been declining significantly. The more negative trend for the thicker multi-year ice area than that for the perennial ice area implies that the average thickness of the ice, and hence the ice volume, has also been declining. Drastic changes in the multi-year ice coverage from QuikScat (satellite radar scatterometer) data, validated using high-resolution Synthetic Aperture Radar data (Kwok, 2004; Nghiem et al., 2007), have also been reported. Some of these changes have been attributed to the near zero replenishment of the Arctic multi-year ice cover by ice that survives the summer (Kwok, 2007).
Observations: Cryosphere Chapter 4
4.2.2.4 Ice Thickness and Volume
For the Arctic, there are several techniques available for estimating the thickness distribution of sea ice. Combined data sets of draft and thickness from submarine sonars, satellite altimetry and airborne elec- tromagnetic sensing provide broadly consistent and strong evidence of decrease in Arctic sea ice thickness in recent years (Figure 4.6c).
Data collected by upward-looking sonar on submarines operating beneath the Arctic pack ice provided the first evidence of ‘basin-wide’ decreases in ice thickness (Wadhams, 1990). Sonar measurements are of average draft (the submerged portion of sea ice), which is converted to thickness by assuming an average density for the measured floe including its snow cover. With the then available submarine records, Rothrock et al. (1999) found that ice draft in the mid-1990s was less than that measured between 1958 and 1977 in each of six regions within the Arctic Basin. The change was least (–0.9 m) in the Beau- fort and Chukchi seas and greatest (–1.7 m) in the Eurasian Basin. The decrease averaged about 42% of the average 1958 to 1977 thickness. This decrease matched the decline measured in the Eurasian Basin between 1976 and 1996 using UK submarine data (Wadhams and Davis, 2000), which was 43%.
A subsequent analysis of US Navy submarine ice draft (Rothrock et al., 2008) used much richer and more geographically extensive data from 34 cruises within a data release area that covered almost 38% of the area of the Arctic Ocean. These cruises were equally distributed in spring and autumn over a 25-year period between 1975 and 2000. Observational uncertainty associated with the ice draft from these is 0.5 m (Rothrock and Wensnahan, 2007). Multiple regression analysis was used to separate the interannual changes (Figure 4.6c), the annual cycle and the spatial distribution of draft in the observations. Results of that analysis show that the annual mean ice thickness declined from a peak of 3.6 m in 1980 to 2.4 m in 2000, a decrease of 1.2 m. Over the period, the most rapid change was –0.08 m yr–1^ in 1990.
The most recent submarine record, Wadhams et al. (2011), found that tracks north of Greenland repeated between the winters of 2004 and 2007 showed a continuing shift towards less multi-year ice.
Satellite altimetry techniques are now capable of mapping sea ice free- board to provide relatively comprehensive pictures of the distribution of Arctic sea ice thickness. Similar to the estimation of sea ice thick- ness from ice draft, satellite measured freeboard (the height of sea ice above the water surface) is converted to thickness, assuming an aver- age density of ice and snow. The principal challenges to accurate thick- ness estimation using satellite altimetry are in the discrimination of ice and open water, and in estimating the thickness of the snow cover.
Since 1993, radar altimeters on the European Space Agency (ESA), European Remote Sensing (ERS) and Envisat satellites have provided Arctic observations south of 81.5°N. With the limited latitudinal reach of these altimeters, however, it has been difficult to infer basin-wide changes in thickness. The ERS-1 estimates of ice thickness show a downward trend but, because of the high variability and short time series (1993–2001), Laxon et al. (2003) concluded that the trend in a region of mixed seasonal and multi-year ice (i.e., below 81.5°N) cannot
be considered as significant. Envisat observations showed a large decrease in thickness (0.25 m) following September 2007 when ice extent was the second lowest on record (Giles et al., 2008b). This was associated with the large retreat of the summer ice cover, with thinning regionally confined to the Beaufort and Chukchi seas, but with no sig- nificant changes in the eastern Arctic. These results are consistent with those from the NASA Ice, Cloud and land Elevation Satellite (ICESat) laser altimeter (see comment on ICESat data in Section 4.4.2.1), which show thinning in the same regions between 2007 and 2008 (Kwok,
- (Figure 4.5). Large decreases in thickness due to the 2007 mini- mum in summer ice are clearly seen in both the radar and laser altim- eter thickness estimates. The coverage of the laser altimeter on ICESat (which ceased opera- tion in 2009) extended to 86°N and provided a more complete spatial pattern of the thickness distribution in the Arctic Basin (Figure 4.6c). Thickness estimates are consistently within 0.5 m of sonar measure- ments from near-coincident submarine tracks and profiles from sonar moorings in the Chukchi and Beaufort seas (Kwok, 2009). Ten ICESat campaigns between autumn 2003 and spring 2008 showed seasonal differences in thickness and thinning and volume losses of the Arctic Ocean ice cover (Kwok, 2009). Over these campaigns, the multi-year sea ice thickness in spring declined by ~0.6 m (Figure 4.5), while the average thickness of the first-year ice (~2 m) had a negligible trend. The average sea ice volume inside the Arctic Basin in spring (February/ March) was ~14,000 km^3. Between 2004 and 2008, the total multi-year ice volume in spring (February/March) experienced a net loss of 6300 km^3 (>40%). Residual differences between sonar mooring and satellite thicknesses suggest basin-scale volume uncertainties of approximate- ly 700 km^3. The rate of volume loss (–1237 km^3 yr–1) during autumn (October/November), while highlighting the large changes during the short ICESat record compares with a more moderate loss rate (– ± 100 km^3 yr–1) over a 31-year period (1979–2010) estimated from a sea ice reanalysis study using the Pan-Arctic Ice-Ocean Modelling and Assimilation system (Schweiger et al., 2011). The CryoSat-2 radar altimeter (launched in 2010), which provides cov- erage up to 89°N, has provided new thickness and volume estimates of Arctic Ocean sea ice (Laxon et al., 2013). These show that the ice volume inside the Arctic Basin decreased by a total of 4291 km^3 in autumn (October/November) and 1479 km^3 in winter (February/March) between the ICESat (2003–2008) and CryoSat-2 (2010–2012) periods. Based on ice thickness estimates from sonar moorings, an inter-satel- lite bias between ICESat and CryoSat-2 of 700 km^3 can be expected. This is much less than the change in volume between the two periods. Airborne electro-magnetic (EM) sounding measures the distance between an EM instrument near the surface or on an aircraft and the ice/water interface, and provides another method to measure ice thick- ness. Uncertainties in these thickness estimates are 0.1 m over level ice. Comparison with drill-hole measurements over a mix of level and ridged ice found differences of 0.17 m (Haas et al., 2011). Repeat EM surveys in the Arctic, though restricted in time and space, have provided a regional view of the changing ice cover. From repeat ground-based and helicopter-borne EM surveys, Haas et al. (2008) found significant thinning in the region of the Transpolar Drift (an
Observations: Cryosphere Chapter 4
Strait. Comparison of volume outflow using ICESat thickness estimates (Spreen et al., 2009) with earlier estimates by Kwok and Rothrock (1999) and Vinje (2001) using thicknesses from moored upward look- ing sonars shows no discernible change.
Between 2005 and 2008, more than a third of the thicker and older sea ice loss occurred by transport of thick, multi-year ice, typically found west of the Canadian Archipelago, into the southern Beaufort Sea, where it melted in summer (Kwok and Cunningham, 2010). Uncertain- ties remain in the relative contributions of in-basin melt and export to observed changes in Arctic ice volume loss, and it has also been shown that export of thicker ice through Nares Strait could account for a small fraction of the loss (Kwok, 2005).
4.2.2.6 Timing of Sea Ice Advance, Retreat and Ice Season Duration; Length of Melt Season
Importantly from both physical and biological perspectives, strong regional changes have occurred in the seasonality of sea ice in both polar regions (Massom and Stammerjohn, 2010; Stammerjohn et al., 2012). However, there are distinct regional differences in when sea- sonally the change is strongest (Stammerjohn et al., 2012).
Seasonality collectively describes the annual time of sea ice advance and retreat, and its duration (the time between day of advance and retreat). Daily satellite ice-concentration records (1979–2012) are used to determine the day to which sea ice advanced, and the day from which it retreated, for each satellite pixel location. Maps of the timing of sea ice advance, retreat and duration are derived from these data (see Parkinson (2002) and Stammerjohn et al. (2008) for detailed methods).
Most regions in the Arctic show trends towards shorter ice season duration. One of the most rapidly changing areas (showing great- er than 2 days yr–1^ change) extends from the East Siberian Sea to the western Beaufort Sea. Here, between 1979 and 2011, sea ice advance occurred 41 ± 6 days later (or 1.3 ± 0.2 days yr–1), sea ice retreat 49 ± 7 days earlier (–1.5 ± 0.2 days yr–1), and duration became 90 ± 16 days shorter (–2.8 ± 0.5 days yr–1) (Stammerjohn et al., 2012). This 3-month lengthening of the summer ice-free season places Arctic summer sea ice extent loss into a seasonal perspective and underscores impacts to the marine ecosystem (e.g., Grebmeier et al., 2010).
The timing of surface melt onset in spring, and freeze-up in autumn, can be derived from satellite microwave data as the emissivity of the surface changes significantly with snow melt (Smith, 1998; Drobot and Anderson, 2001; Belchansky et al., 2004). The amount of solar energy absorbed by the ice cover increases with the length of the melt season. Longer melt seasons with lower albedo surfaces (wet snow, melt ponds and open water) increase absorption of incoming shortwave radiation and ice melt (Perovich et al., 2007). Hudson (2011) estimates that the observed reduction in Arctic sea ice has contributed approximately 0.1 W m–2^ of additional global radiative forcing, and that an ice-free summer Arctic Ocean will result in a forcing of about 0.3 W m–2. The satellite record (Markus et al., 2009) shows a trend toward earlier melt and later freeze-up nearly everywhere in the Arctic (Figure 4.6e). Over the last 34 years, the mean melt season over the Arctic ice cover has
increased at a rate of 5.7 ± 0.9 days per decade. The largest and most significant trends (at the 99% level) of more than 10 days per decade are seen in the coastal margins and peripheral seas: Hudson Bay, the East Greenland Sea, the Laptev/East Siberian seas, and the Chukchi/ Beaufort seas. 4.2.2.7 Arctic Polynyas High sea ice production in coastal polynyas (anomalous regions of open water or low ice concentration) over the continental shelves of the Arctic Ocean is responsible for the formation of cold saline water, which contributes to the maintenance of the Arctic Ocean halocline (see Glossary). A new passive microwave algorithm has been used to estimate thin sea ice thicknesses (<0.15 m) in the Arctic Ocean (Tamura and Ohshima, 2011), providing the first circumpolar mapping of sea ice production in coastal polynyas. High sea ice production is confined to the most persistent Arctic coastal polynyas, with the highest ice pro- duction rate being in the North Water Polynya. The mean annual sea ice production in the 10 major Arctic polynyas is estimated to be 2942 ± 373 km^3 and decreased by 462 km^3 between 1992 and 2007 (Tamura and Ohshima, 2011). 4.2.2.8 Arctic Land-Fast Ice Shore- or land-fast ice is sea ice attached to the coast. Land-fast ice along the Arctic coast is usually grounded in shallow water, with the seaward edge typically around the 20 to 30 m isobath (Mahoney et al., 2007). In fjords and confined bays, land-fast ice extends into deeper water. There are no reliable estimates of the total area or interannual variabil- ity of land-fast ice in the Arctic. However, both significant and non-sig- nificant trends have been observed regionally. Long-term monitoring near Hopen, Svalbard, revealed thinning of land-fast ice in the Barents Sea region by 11 cm per decade between 1966 and 2007 (Gerland et al., 2008). Between 1936 and 2000, the trends in land-fast ice thick- ness (in May) at four Siberian sites (Kara Sea, Laptev Sea, East Siberi- an Sea, Chukchi Sea) are insignificant (Polyakov et al., 2003). A more recent composite time series of land-fast ice thickness between the mid 1960s and early 2000s from 15 stations along the Siberian coast revealed an average rate of thinning of 0.33 cm yr–1(Polyakov et al., 2010). End-of-winter ice thickness for three stations in the Canadian Arctic reveal a small downward trend at Eureka, a small positive trend at Resolute Bay, and a negligible trend at Cambridge Bay (updated from Brown and Coté, 1992; Melling, 2012), but these trends are small and not statistically significant. Even though the trend in the land-fast ice extent near Barrow, Alaska has not been significant (Mahoney et al., 2007), relatively recent observations by Mahoney et al. (2007) and Druckenmiller et al. (2009) found longer ice-free seasons and thinner land-fast ice compared to earlier records (Weeks and Gow, 1978; Barry et al., 1979). As freeze-up happens later, the growth season shortens and the thinner ice breaks up and melts earlier. 4.2.2.9 Decadal Trends in Arctic Sea Ice The average decadal extent of Arctic sea ice has decreased in every season and in every successive decade since satellite observations
Chapter 4 Observations: Cryosphere
commenced. The data set is robust with continuous and consistent global coverage on a daily basis thereby providing very reliable trend results (very high confidence). The annual Arctic sea ice cover very likely declined within the range 3.5 to 4.1% per decade (0.45 to 0. million km^2 per decade) during the period 1979–2012 with larger changes occurring in summer and autumn (very high confidence). Much larger changes apply to the perennial ice (the summer minimum extent) which very likely decreased in the range from 9.4 % to 13.6 % per decade (0.73 to 1.07 million km^2 per decade) and multiyear sea ice (more than 2 years old) which very likely declined in the range from 11.0 % to 16.0% per decade (0.66 to 0.98 million km^2 per decade) (very high confidence; Figure 4.4b). The rate of decrease in ice area has been greater than that in extent (Figure 4.4b) because the ice con- centration has also decreased. The decline in multiyear ice cover as observed by QuikScat from 1992 to 1910 is presented in Figure 4.6b and shown to be consistent with passive microwave data (Figure 4.4b). The decrease in perennial and multi-year ice coverage has resulted in a strong decrease in ice thickness, and hence in ice volume. Declassified submarine sonar measurements, covering ~38% of the Arctic Ocean, indicate an overall mean winter thickness of 3.64 m in 1980, which likely decreased by 1.8 [1.3 to 2.3] m by 2008 (high confidence, Figure 4.6c). Between 1975 and 2000, the steepest rate of decrease was 0. m yr–1^ in 1990 compared to a slightly higher winter/summer rate of 0.10/0.20 m yr–1^ in the 5-year ICESat record (2003–2008). This com- bined analysis (Figure 4.6c) shows a long-term trend of sea ice thinning that spans five decades. Satellite measurements made in the period 2010–2012 show a decrease in basin-scale sea ice volume compared to those made over the period 2003–2008 (medium confidence). The Arctic sea ice is becoming increasingly seasonal with thinner ice, and it will take several years for any recovery. The decreases in both concentration and thickness reduces sea ice strength reducing its resistance to wind forcing, and drift speed has increased (Figure 4.6d) (Rampal et al., 2009; Spreen et al., 2011). Other significant changes to the Arctic Ocean sea ice include lengthening in the duration of the surface melt on perennial ice of 6 days per decade (Figure 4.6e) and a nearly 3-month lengthening of the ice-free season in the region from the East Siberian Sea to the western Beaufort Sea. 4.2.3 Antarctic Sea Ice The Antarctic sea ice cover is largely seasonal, with average extent var- ying from a minimum of about 3 × 10^6 km^2 in February to a maximum of about 18 × 10^6 km^2 in September (Zwally et al., 2002a; Comiso et al., 2011). The relatively small fraction of Antarctic sea ice that survives the summer is found mostly in the Weddell Sea, but with some perennial ice also surviving on the western side of the Antarctic Peninsula and in small patches around the coast. As well as being mostly first-year ice, Antarctic sea ice is also on average thinner, warmer, more saline and more mobile than Arctic ice (Wadhams and Comiso, 1992). These characteristics, which reduce the capabilities of some remote sensing techniques, together with its more distant location from inhabited con- tinents, result in far less being known about the properties of Antarctic sea ice than of that in the Arctic.
4.2.3.1 Total Antarctic Sea Ice Extent and Concentration Figure 4.7a shows the seasonal variability of Antarctic sea ice extent using 34 years of satellite passive microwave data updated from Comiso and Nishio (2008). In contrast to the Arctic, decadal monthly averages almost overlap with each other, and the seasonal variability of the total Antarctic sea ice cover has not changed much over the period. In winter, the values for the 1999–2008 decade were slightly higher than those of the other decades; whereas in autumn the values for 1989–1998 and 1999–2008 decades were higher than those of 1979–1988. There was more seasonal variability in the period 2009– 2012 than for earlier decadal periods, with relatively high values in late autumn, winter and spring. Trend maps for winter, spring, summer and autumn extent are present- ed in Figure 4.7 (b, c, d and e respectively). The seasonal trends are sig- nificant mainly near the ice edge, with the values alternating between positive and negative around Antarctica. Such an alternating pattern is similar to that described previously as the Antarctic Circumpolar Wave (ACW) (White and Peterson, 1996) but the ACW may not be associated with the trends because the trends have been strongly positive in the Ross Sea and negative in the Bellingshausen/Amundsen seas but with almost no trend in the other regions (Comiso et al., 2011). In the winter, negative trends are evident at the tip of the Antarctic Peninsula and the western part of the Weddell Sea, while positive trends are prev- alent in the Ross Sea. The patterns in spring are very similar to those of winter, whereas in summer and autumn negative trends are mainly confined to the Bellingshausen/Amundsen seas, while positive trends are dominant in the Ross Sea and the Weddell Sea. The regression trend in the monthly anomalies of Antarctic sea ice extent from November 1978 to December 2012 (updated from Comiso and Nishio, 2008) is slightly positive, at 1.5 ± 0.3% per decade, or 0. to 0.20 million km^2 per decade (very likely) (see FAQ 4.1). The seasonal trends in ice extent are 1.2 ± 0.5%, 1.0 ± 0.5%, 2.5 ± 2.0% and 3. ± 2.0% per decade (very likely) in winter, spring, summer and autumn, respectively, as updated from Comiso et al. (2011). The corresponding trends in ice area (also updated) are 1.9 ± 0.7%, 1.6 ± 0.5%, 3.0 ± 2.1%, and 4.4 ± 2.3% per decade (very likely). The values are all pos- itive, with the largest trends occurring in the autumn. The trends are consistently higher for ice area than ice extent, indicating less open water (possibly due to less storms and divergence) within the pack in later years. Trends reported by Parkinson and Cavalieri (2012) using data from 1978 to 2010 are slightly different, in part because they cover a different time period (see Supplementary Material). The overall interannual trends for various sectors around Antarctica are given in FAQ 4.1, and show large regional variability. Changes in ice drift and wind patterns as reported by Holland and Kwok (2012) may be related to this phenomenon. 4.2.3.2 Antarctic Sea Ice Thickness and Volume Since AR4, some advances have been made in determining the thick- ness of Antarctic sea ice, particularly in the use of ship-based obser- vations and satellite altimetry. However, there is still no information on large-scale Antarctic ice thickness change. Worby et al. (2008) compiled 25 years of ship-based data from 83 Antarctic voyages on
Chapter 4 Observations: Cryosphere
Figure 4.7 | (a) Plots of decadal averages of daily sea ice extent in the Antarctic (1979–1988 in red, 1989–1998 in blue, 1999– 2008 in gold) and a 4-year average daily ice extent from 2009 to 2012 in black. Maps indicate ice concentration trends (1979–2012) in (b) winter, (c) spring, (d) summer and (e) autumn (updated from Comiso, 2010).
0
5
15
20
10
a) Daily ice extent
J F M A M J J A S O N D
Ice extent (
6 km
2 )
1979- 1989- 1999- 2009-
b) Winter (JJA) c) Spring (SON)
d) Summer (DJF) e) Autumn (MAM)
135 o^ W
45 o^ E
60 o^ S 50 o^ S
-2.4 -1.6 -0.8 0.0 0.8 1.6 2. Trend (% IC yr-1)
which routine observations of sea ice and snow properties were made. Their compilation included a gridded data set that reflects the regional differences in sea ice thickness. A subset of these ship observations, and ice charts, was used by DeLiberty et al. (2011) to estimate the annual cycle of sea ice thickness and volume in the Ross Sea, and to investigate the relationship between ice thickness and extent. They found that maximum sea ice volume was reached later than maxi- mum extent. While ice is advected to the northern edge and melts, the interior of the sea ice zone is supplied with ice from higher latitudes and continues to thicken by thermodynamic growth and deformation. Satellite retrievals of sea ice freeboard and thickness in the Antarctic (Mahoney et al., 2007; Zwally et al., 2008; Xie et al., 2011) are under
development but progress is limited by knowledge of snow thickness and the paucity of suitable validation data sets. A recent analysis of the ICESat record by Kurtz and Markus (2012), assuming zero ice freeboard, found negligible trends in ice thickness over the 5-year record. 4.2.3.3 Antarctic Sea Ice Drift Using a 19-year data set (1992–2010) of satellite-tracked sea ice motion, Holland and Kwok (2012) found large and statistically sig- nificant decadal trends in Antarctic ice drift that in most sectors are caused by changes in local winds. These trends suggest acceleration of the wind-driven Ross Gyre and deceleration of the Weddell Gyre. The changes in meridional ice transport affect the freshwater budget near the Antarctic coast. This is consistent with the increase of 30,000 km^2 yr–1^ in the net area export of sea ice from the Ross Sea shelf coastal polynya region between 1992 and 2008 (Comiso et al., 2011). Assum- ing an annual average thickness of 0.6 m, Comiso et al. (2011) estimat- ed an increase in volume export of 20 km^3 yr–1^ which is similar to the rate of production in the Ross Sea coastal polynya region for the same period discussed in Section 4.2.3.5. 4.2.3.4 Timing of Sea Ice Advance, Retreat and Ice Season Duration In the Antarctic there are regionally different patterns of strong change in ice duration (>2 days yr–1). In the northeast and west Antarctic Peninsula and southern Bellingshausen Sea region, later ice advance (+61 ± 15 days), earlier retreat (–39 ± 13 days) and shorter duration (+100 ± 31 days, a trend of –3.1 ± 1.0 days yr–1) occurred over the period 1979/1980–2010/2011 (Stammerjohn et al., 2012). These changes have strong impacts on the marine ecosystem (Montes- Hugo et al., 2009; Ducklow et al., 2011). The opposite is true in the adjacent western Ross Sea, where substantial lengthening of the ice season of 79 ± 12 days has occurred (+2.5 ± 0.4 days yr–1) due to earlier advance (+42 ± 8 days) and later retreat (–37 ± 8 days). Patterns of change in the relatively narrow East Antarctic sector are generally of a lower magnitude and zonally complex, but in certain regions involve changes in the timing of sea ice advance and retreat of the order of ±1 to 2 days yr–1^ (for the period 1979–2009) (Massom et al., 2013). 4.2.3.5 Antarctic Polynyas Polynyas are commonly found along the coast of Antarctica. There are two different processes that cause a polynya. Warm water upwelling keeps the surface water near the freezing point and reduces ice pro- duction (sensible heat polynya), and wind or ocean currents move ice away and increase further ice production (latent heat polynya). An increase in the extent of coastal polynyas in the Ross Sea caused increased ice production (latent heat effect) that is primarily respon- sible for the positive trend in ice extent in the Antarctic (Comiso et al., 2011). Drucker et al. (2011) show that in the Ross Sea, the net ice export equals the annual ice production in the Ross Sea polynya (approximately 400 km^3 in 1992), and that ice production increased by 20 km^3 yr–1^ from 1992 to 2008. However, the ice production in the Wed- dell Sea, which is three times less, has had no statistically significant
Observations: Cryosphere Chapter 4
Frequently Asked Questions
FAQ 4.1 | How Is Sea Ice Changing in the Arctic and Antarctic?
The sea ice covers on the Arctic Ocean and on the Southern Ocean around Antarctica have quite different charac- teristics, and are showing different changes with time. Over the past 34 years (1979–2012), there has been a down- ward trend of 3.8% per decade in the annual average extent of sea ice in the Arctic. The average winter thickness of Arctic Ocean sea ice has thinned by approximately 1.8 m between 1978 and 2008, and the total volume (mass) of Arctic sea ice has decreased at all times of year. The more rapid decrease in the extent of sea ice at the summer minimum is a consequence of these trends. In contrast, over the same 34-year period, the total extent of Antarctic sea ice shows a small increase of 1.5% per decade, but there are strong regional differences in the changes around the Antarctic. Measurements of Antarctic sea ice thickness are too few to be able to judge whether its total volume (mass) is decreasing, steady, or increasing. A large part of the total Arctic sea ice cover lies above 60°N (FAQ 4.1, Figure 1) and is surrounded by land to the south with openings to the Canadian Arctic Archipelago, and the Bering, Barents and Greenland seas. Some of the ice within the Arctic Basin survives for several seasons, growing in thickness by freezing of seawater at the base and by deformation (ridging and rafting). Seasonal sea ice grows to only ~2 m in thickness but sea ice that is more than 1 year old (perennial ice) can be several metres thicker. Arctic sea ice drifts within the basin, driven by wind and ocean currents: the mean drift pattern is dominated by a clockwise circulation pattern in the western Arctic and a Transpolar Drift Stream that transports Siberian sea ice across the Arctic and exports it from the basin through the Fram Strait. Satellites with the capability to distinguish ice and open water have provided a picture of the sea ice cover changes. Since 1979, the annual average extent of ice in the Arctic has decreased by 3.8% per decade. The decline in extent at the end of summer (in late September) has been even greater at 11% per decade, reaching a record minimum in
- The decadal average extent of the September minimum Arctic ice cover has decreased for each decade since satellite records began. Submarine and satellite records suggest that the thickness of Arctic ice, and hence the total volume, is also decreasing. Changes in the relative amounts of perennial and seasonal ice are contributing to the reduction in ice volume. Over the 34-year record, approximately 17% of this type of sea ice per decade has been lost to melt and export out of the basin since 1979 and 40% since 1999. Although the area of Arctic sea ice coverage can fluctuate from year to year because of variable seasonal production, the proportion of thick perennial ice, and the total sea ice volume, can recover only slowly. Unlike the Arctic, the sea ice cover around Antarctica is constrained to latitudes north of 78°S because of the pres- ence of the continental land mass. The Antarctic sea ice cover is largely seasonal, with an average thickness of only ~1 m at the time of maximum extent in September. Only a small fraction of the ice cover survives the summer minimum in February, and very little Antarctic sea ice is more than 2 years old. The ice edge is exposed to the open ocean and the snowfall rate over Antarctic sea ice is higher than in the Arctic. When the snow load from snowfall is sufficient to depress the ice surface below sea level, seawater infiltrates the base of the snow pack and snow-ice is formed when the resultant slush freezes. Consequently, snow-to-ice conversion (as well as basal freezing as in the Arctic) contributes to the seasonal growth in ice thickness and total ice volume in the Antarctic. Snow-ice forma- tion is sensitive to changes in precipitation and thus changes in regional climate. The consequence of changes in precipitation on Antarctic sea ice thickness and volume remains a focus for research. Unconstrained by land boundaries, the latitudinal extent of the Antarctic sea ice cover is highly variable. Near the Antarctic coast, sea ice drift is predominantly from east to west, but further north, it is from west to east and highly divergent. Distinct clockwise circulation patterns that transport ice northward can be found in the Weddell and Ross seas, while the circulation is more variable around East Antarctica. The northward extent of the sea ice cover is controlled in part by the divergent drift that is conducive in winter months to new ice formation in persistent open water areas (polynyas) along the coastlines. These zones of ice formation result in saltier and thus denser ocean water and become one of the primary sources of the deepest water found in the global oceans. Over the same 34-year satellite record, the annual extent of sea ice in the Antarctic increased at about 1.5% per decade. However, there are regional differences in trends, with decreases seen in the Bellingshausen and Amundsen seas, but a larger increase in sea ice extent in the Ross Sea that dominates the overall trend. Whether the smaller overall increase in Antarctic sea ice extent is meaningful as an indicator of climate is uncertain because the extent (continued on next page)
Observations: Cryosphere Chapter 4
waves, and strong wind events that cause the ice to break-up. Histor- ical records of Antarctic land-fast ice extent, such as that of Kozlovsky et al. (1977) covering 0° to 160°E, were limited by sparse and sporad- ic sampling. Recently, using cloud-free Moderate Resolution Imaging Spectrometer (MODIS) composite images, Fraser et al. (2012) derived a high-resolution time series of land-fast sea ice extent along the East Antarctic coast, showing a statistically significant increase (1. ± 0.30% yr–1) between March 2000 and December 2008. There is a strong increase in the Indian Ocean sector (20°E to 90°E, 4.07 ± 0.42% yr–1), and a non-significant decrease in the sector from 90°E to 160°E (–0.40 ± 0.37% yr–1). An apparent shift from a negative to a positive trend was noted in the Indian Ocean sector from 2004, which coincid- ed with greater interannual variability. Although significant changes are observed, this record is only 9 years in length.
4.2.3.7 Decadal Trends in Antarctic Sea Ice
For the Antarctic, any changes in many sea ice characteristics are unknown. There has been a small but significant increase in total annual mean sea ice extent that is very likely in the range of 1.2 to 1.8 % per decade between 1979 and 2012 (0.13 to 0.20 million km^2 per decade) (very high confidence). There was also a greater increase in ice area associated with an increase in ice concentration. But there are strong regional differences within this total, with some regions increasing in extent/area and some decreasing (high confidence). Similarly, there are contrasting regions around the Antarctic where the ice-free season has lengthened, and others where it has decreased over the satellite period (high confidence). There are still inadequate data to make any assess- ment of changes to Antarctic sea ice thickness and volume.
4.3 Glaciers
This section considers all perennial surface land ice masses (defined in 4.1 and Glossary) outside of the Antarctic and Greenland ice sheets. Glaciers occur where climate conditions and topographic characteristics allow snow to accumulate over several years and to transform grad- ually into firn (snow that persists for at least one year) and finally to ice. Under the force of gravity, this ice flows downwards to elevations with higher temperatures where various processes of ablation (loss of snow and ice) dominate over accumulation (gain of snow and ice). The sum of all accumulation and ablation processes determines the mass balance of a glacier. Accumulation is in most regions due mainly to solid precipitation (in general snow), but also results from refreezing of liquid water, especially in polar regions or at high altitudes where firn remains below melting temperature. Ablation is, in most regions, mainly due to surface melting with subsequent runoff, but loss of ice by calving (on land or in water; see Glossary) or sublimation (important in dry regions) can also dominate. Re-distribution of snow by wind and avalanches can contribute to both accumulation and ablation. The energy and mass fluxes governing the surface mass balance are direct- ly linked to atmospheric conditions and are modified by topography (e.g., due to shading). Glaciers are sensitive climate indicators because they adjust their size in response to changes in climate (e.g., tempera- ture and precipitation) (FAQ 4.2). Glaciers are also important season- al to long-term hydrologic reservoirs (WGII, Chapter 3) on a regional scale and a major contributor to sea level rise on a global scale (see
Section 4.3.3.4 and Chapter 13). In the following, we report global glacier coverage (Section 4.3.1), how changes in length, area, volume and mass are determined (Section 4.3.2) and the observed changes in these parameters through time (Section 4.3.3). 4.3.1 Current Area and Volume of Glaciers The total area covered by glaciers was only roughly known in AR4, resulting in large uncertainties for all related calculations (e.g., overall glacier volume or mass changes). Since AR4, the world glacier invento- ry (WGMS, 1989) was gradually extended by Cogley (2009a) and Radić and Hock (2010); and for AR5, a new globally complete data set of gla- cier outlines (Randolph Glacier Inventory (RGI)) was compiled from a wide range of data sources from the 1950s to 2010 with varying levels of detail and quality (Arendt et al., 2012). Regional glacier-covered areas for 19 regions were extracted from the RGI and supplemented with the percentage of the area covered by glaciers terminating in tide- water (Figure 4.8 and Table 4.2). The areas covered by glaciers that are in contact with freshwater lakes are only locally available. The separa- tion of so-called peripheral glaciers from the ice sheets in Greenland and Antarctica is not easy. A new detailed inventory of the glaciers in Greenland (Rastner et al., 2012) allows for estimation of their area, volume, and mass balance separately from those of the ice sheet. This separation is still incomplete for Antarctica, and values discussed here (Figures 4.1, 4.8 to 4.11, Tables 4.2 and 4.4) refer to the glaciers on the islands in the Antarctic and Sub-Antarctic (Bliss et al., 2013) but exclude glaciers on the mainland of Antarctica that are separate from the ice sheet. Regionally variable accuracy of the glacier outlines leads to poorly quantified uncertainties. These uncertainties, along with the regional variation in the minimum size of glaciers included in the inventory, and the subdivision of contiguous ice masses, also makes the total number of glaciers uncertain; the current best estimate is around 170,000 covering a total area of about 730,000 km^2. When summed up, nearly 80% of the glacier area found in regions Antarctic and Subantarctic (region 19), Canadian Arctic (regions 3 and 4), High Mountain Asia (regions 13, 14 and 15), Alaska (region 5), and Green- land (region 17) (Table 4.2). From the glacier areas in the new inventory, total glacier volumes and masses have been determined by applying both simple scaling relations and ice-dynamical considerations (Table 4.2, and references therein), however, both methods are calibrated with only a few hun- dred glacier thickness measurements. This small sample means that uncertainties are large and difficult to quantify. The range of values as derived from four global-scale studies for each of the 19 RGI regions is given in Table 4.2, suggesting a global glacier mass that is likely between 114,000 and 192,000 Gt (314 to 529 mm SLE). The numbers and areas of glaciers reported in Table 4.2 are directly taken from RGI 2.0 (Arendt et al., 2012), with updates for the Low Latitudes (region 16) and the Southern Andes (region 17). 4.3.2 Methods to Measure Changes in Glacier Length, Area and Volume/Mass To measure changes in glacier length, area, mass and volume, a wide range of observational techniques has been developed. Each technique has individual benefits over specific spatial and temporal scales; their
Chapter 4 Observations: Cryosphere
Region Region Name Number of Glaciers ( km^ Area (^2) )^ Percent of total area fraction (%)^ Tidewater Mass (minimum) (Gt)
Mass (maximum) (Gt)
Mean SLE (mm) 1 Alaska 23,112 89,267 12.3 13.7 16,168 28,021 54. 2 Western Canada and USA 15,073 14,503.5 2.0 0 906 1148 2. 3 Arctic Canada North 3318 103,990.2 14.3 46.5 22,366 37,555 84. 4 Arctic Canada South 7342 40,600.7 5.6 7.3 5510 8845 19. 5 Greenland 13,880 87,125.9 12.0 34.9 10,005 17,146 38. 6 Iceland 290 10,988.6 1.5 0 2390 4640 9. 7 Svalbard 1615 33,672.9 4.6 43.8 4821 8700 19. 8 Scandinavia 1799 2833.7 0.4 0 182 290 0. 9 Russian Arctic 331 51,160.5 7.0 64.7 11,016 21,315 41. 10 North Asiaa^4403 3425.6 0.4 0 109 247 0. 11 Central Europe 3920 2058.1 0.3 0 109 125 0. 12 Caucasus 1339 1125.6 0.2 0 61 72 0. 13 Central Asia 30,200 64,497 8.9 0 4531 8591 16. 14 South Asia (West) 22,822 33,862 4.7 0 2900 3444 9. 15 South Asia (East) 14,006 21,803.2 3.0 0 1196 1623 3. 16 Low Latitudesa^2601 2554.7 0.6 0 109 218 0. 17 Southern Andesa^ 15,994 29,361.2 4.5 23.8 4241 6018 13. 18 New Zealand 3012 1160.5 0.2 0 71 109 0. 19 Antarctic and Sub-Antarctic 3274 13,2267.4 18.2 97.8 27,224 43,772 96. Total 168,331 726,258.3 38.5 113,915 191,879 412.
Table 4.2 | The 19 regions used throughout this chapter and their respective glacier numbers and area (absolute and in percent) are derived from the RGI 2.0 (Arendt et al., 2012); the tidewater fraction is from Gardner et al. (2013). The minimum and maximum values of glacier mass are the minimum and maximum of the estimates given in four studies: Grinsted (2013), Huss and Farinotti (2012), Marzeion et al. (2012) and Radić et al. (2013). The mean sea level equivalent (SLE) of the mean glacier mass is the mean of estimates from the same four studies, using an ocean area of 362.5 × 10^6 km^2 for conversion. All values were derived with globally consistent methods; deviations from more precise national data sets are thus possible. Ongoing improvements may lead to revisions of these (RGI 2.0) numbers in future releases of the RGI.
Notes: a For regions 10, 16 and 17 the number and area of glaciers are corrected to allow for over-inclusion of seasonal snow in the glacierized extent of RGI 2.0 and for improved outlines (region 10) compared to RGI 2.0 (updated from, Arendt et al., 2012).
main characteristics are summarized in Table 4.3. Monitoring programs include complex climate-related observations at a few glaciers, index measurements of mass balance at about a hundred glaciers, annual length changes for a few hundred glaciers, and repeat geodetic esti- mates of area and volume changes at regional scales using remote sensing methods (e.g., Haeberli et al., 2007). Although in situ meas- urements of glacier changes are biased towards glaciers that are easily accessible, comparatively small and simple to interpret, a large propor- tion of all glaciers in the world is debris covered or tidewater calving (see Table 4.2) and changes of such glaciers are more difficult to inter- pret in climatic terms (Yde and Pasche, 2010). In addition, many of the remote-sensing based assessments do not discriminate these types. 4.3.2.1 Length Change Measurements For the approximately 500 glaciers worldwide that are regularly observed, front variations (commonly called length changes) are usu- ally obtained through annual measurements of the glacier terminus position. Globally coordinated observations were started in 1894, providing one of the longest available time series of environmental change (WGMS, 2008). More recently, particularly in regions that are
difficult to access, aerial photography and satellite imaging have been used to determine glacier length changes over the past decades. For selected glaciers globally, historic terminus positions have been recon- structed from maps, photographs, satellite imagery, also paintings, dated moraines and other sources (e.g., Masiokas et al., 2009; Lopez et al., 2010; Nussbaumer et al., 2011; Davies and Glasser, 2012; Leclercq and Oerlemans, 2012; Rabatel et al., 2013). Early reconstructions are sparsely distributed in both space and time, generally at intervals of decades. The terminus fluctuations of some individual glaciers have been reconstructed for periods of more than 3000 years (Holzhauser et al., 2005), with a much larger number of records available as far back as the 16th or 17th centuries (Zemp et al., 2011, and references therein). The reconstructed glacier length records are globally well dis- tributed and were used, for example, to determine the contribution of glaciers to global sea level rise (Leclercq et al., 2011) (Section 4.3.3.4), and for an independent temperature reconstruction at a hemispheric scale (Leclercq and Oerlemans, 2012).
