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2.7 Recent changes in the Antarctic

2.7.1 T rends in the surface observations

The relatively short length of many of the series of observations from the Antarctic research stations makes the assessment of trends in the meteorological conditions across the continent very difficult and it is not usually possible to find trends that are statistically significant. However, an examination of the longest series of data is instructive and in Figure 2.7.1.1 the series of annual mean surface air temperature values from a number of stations are shown, along with trend lines determined from least–squares regression. These stations are located across the continent from the South Pole (Amundsen–Scott), to the high interior plateau (Vostok), coastal stations (Mirny and Halley) to Orcadas in the South Orkney Islands. It can be seen that all the stations show a high degree of inter–annual variability, which is a feature of stations in both polar regions, but that the variability is largest on the western side of the Antarctic Peninsula at Faraday/Vernadsky Station. This station is located close to the northern limit of the Bellingshausen Sea sea ice and small variations in the ice extent are amplified into much larger surface temperature variations depending on whether the ocean west of the station is ice–covered or ice–free in a particular winter. Most of the stations show a small warming trend; the exception being Amundsen–Scott Station, where there has been a slight cooling since the late 1950s. The warming on the western side of the Antarctic Peninsula is larger than elsewhere in the Antarctic and even though the record is not long by the standards of stations outside the Antarctic, the warming trend is statistically significant at the 99% level (King, 1994). In parallel with the warming trend there has also been a statistically significant increase in the number of precipitation reports at the stations on the western side of the Peninsula (Turner et al., 1997). At the moment we do not know whether the climatic changes observed in this area are the result of local factors or because of broader scale circulation changes across the Pacific region. However, there is no evidence that the warming is taking place because of any “global warming” associated with the increased anthropogenic emission of greenhouse gases.

Figure 2.7.1.1 Time series of annual mean surface air temperature at a number of

Antarctic stations. (Data are from Jones and Limbert (1987), updated from recent reports in some cases.)

Antarctic–wide temperature trends were considered in an earlier investigation (Raper et al., 1984) within which seasonal and annual average temperatures for Antarctica were calculated by computing areally–weighted means of all available station data. The annual mean Antarctic temperature showed a warming trend of 0.029ºC per year for the period 1957‑82, which was significant at the 95% level. However, the greatest contribution from this temperature rise came from the stations on the western side of the Antarctic Peninsula.

2.7.2 Ozone over Antarctica

Ozone is one of the key radiatively–active gasses in the stratosphere. Where there is less ozone the stratosphere tends to be colder and at any one location there is a good correlation between the 100 or 70–hPa temperature and the total column ozone. Ozone also absorbs solar uv light and where there is less ozone more uv will reach the surface, posing risks to human health.

The total column ozone was first measured in Antarctica during the IGY (1957–58). The instrument used to make the measurements was the Dobson ozone spectrophotometer, designed by Professor G. M. B. Dobson in the 1920s and still the standard instrument today. A typical value for the total ozone column is around 300 Dobson Units (DU), or 300 milli‑atmosphere–centimetres, which corresponds to a layer of ozone 3 mm thick at the surface. This 3 mm is in reality spread through a column, the bulk of which lies between the tropopause and 40 km altitude, with a maximum at around 17 km altitude.

The seasonal pattern of the total ozone column in Antarctica is linked to the development and breakdown of the winter circumpolar vortex. Historically ozone values were around 300 DU at the beginning of the winter, and similar at the end. During the winter, ozone amounts build up in a circumpolar belt just outside the vortex, due to transport of ozone from source regions in the tropics. Very high values, occasionally exceeding 500 DU may be seen in this belt, which typically lies between about 60° S and 40° S. When the polar vortex breaks down in the spring, this belt of high ozone column air sweeps across Antarctica, typically reaching the area around the central Pacific sector first and the Atlantic sector last. Peak values used to exceed 450 DU, and then slowly declined back to 300 DU during the summer and autumn.

Since the mid 1970s an increasingly different pattern of behaviour is seen – the Antarctic Ozone Hole. At the end of the winter, values are around 10% lower than they were in the 1970s and drop at around 1% per day to reach around 100 DU at the end of September. Values then slowly begin to recover, but the spring warming in the stratosphere is often delayed to the end of November or into December. During this spring period Antarctic personnel should be advised to use high factor blocking creams, even during cloudy weather as the increase in surface uv radiation can lead to severe sun–burn. Peak values in the spring warming are now substantially lower than in pre–Ozone Hole times. Figure 2.7.2.1 shows the seasonal variation of total ozone at Halley for the period 1957 to 1998 and the decline in late spring and early summer values since about the early seventies is very evident as is the presence of very low October values since about the 1980s. Figure 2.7.2.2 shows the total October ozone at Halley since records began: the decrease from about 1975, and certainly from 1980 onwards, is very marked.

The mechanism that controls the development of the Antarctic ozone hole is linked to the dynamics of the winter polar vortex. During the winter, lower stratospheric temperatures drop below –80°C and at these temperatures stratospheric clouds can form. (Figure 2.7.2.3). Observers at stations along the Antarctic Peninsula regularly see these during the late winter as nacreous or “Mother of Pearl” clouds, created by lee–waves off the mountains of the Peninsula. More southerly stations sometimes report them as “ultra–cirrus”, which may cover the entire sky in a faint milky veil. Occasionally there are reports of clouds that resemble noctilucent clouds and it is just possible that such mesospheric clouds are seen in the Antarctic winter; these reports need further investigation. Once the clouds have formed, chemical reactions take place on the cloud surfaces that lock up nitrogen oxides and water and liberate chlorine and bromine (from chloro–flouro-carbons (CFCs), halons and other similar chemicals) in an active form. When exposed to sunlight, photochemical catalytic reactions take place that rapidly destroy ozone.

Figure 2.7.2.1 The seasonal variation of total ozone at Halley for the period 1957–98.

Figure 2.7.2.2 Mean total ozone for

October as measured at Halley since the

1950s.

When the circumpolar vortex is at its strongest the ozone hole tends to be roughly circular, (see, for example, Figure 2.7.2.4) but as it weakens the vortex often becomes strongly elliptical (Figure 2.7.2.5) and often offset from the pole towards the Atlantic. At these times the northern edge of the vortex can reach as far north as 50° S, posing a risk of increased uv exposure for the inhabitants of southern South America and the sub–Antarctic islands. The wave rotation period of the vortex is typically around a month and this can be used to give rough forecasts of the period when such areas are most at risk. Further guidance can be given from the forecast 100 or 70 hPa temperature fields: the –70ºC temperature on the 100‑hPa‑surface gives a rough indication of the edge of the vortex and the –65ºC contour can be used for conservative forecasting.

The mean 100–hPa temperature during October to December is now significantly lower than it was 30 years ago. This is due to a combination of the delayed spring warming and the lower amount of ozone in the stratosphere. (Figure 2.7.2.6 is a comparison of recent measurements of the total ozone at Halley compared to 1957–72 values). This potentially has implications for the surface climate, though the coupling between stratosphere and troposphere is weak.

The size and depth of the ozone hole at its maximum extent now seem to be near their peak. The size of the hole is constrained by the circumpolar vortex, which in turn is constrained by the size of the Antarctic continent. The depth of the hole is constrained by the height range where temperatures are cold enough for the stratospheric clouds to form. We should see a very slow return to the pre–1970 situation, provided that the Montreal Protocol is adhered to, and that there are no other changes to the atmosphere. Greenhouse warming is one such change, and although this warms the lower troposphere it cools the lower stratosphere, making the occurrence of stratospheric clouds more likely. This may delay the recovery of the Ozone Hole past the middle of the 21st century.

Figure 2.7.2.3 Stratospheric clouds viewed from Faraday.

2.7.3 Recent changes to Antarctic ice shelves

2.7.3.1 On the relationship between polar ice sheets, climate and sea–level

In a recent summary of the relationship between polar ice sheets, climate and sea–level rise the Antarctic Co–operative Research Centre, Hobart, Australia issued a position statement on 10 February 2000 (see: http://www.acecrc.org.au/) as follows:

"At the end of the last ice age the melting of the ice–sheets on North America and Europe caused sea level to rise by 120 m. (The last great ice–age had its peak about 18 thousand years ago and retreated to roughly the present situation by about 11 thousand years ago. It is thought that the process was triggered by slight changes in the orbit and inclination of the Earth, which in turn altered the ratio of summer to winter solar heating at high latitudes).

Figure 2.7.2.4 Total ozone for 23 September 1999: an example of the circular nature of the ozone hole when at its greatest depth.

Figure 2.7.2.5 Total ozone for 21 October 1998: an example of the elliptical nature of the ozone hole during the weakening phase (compare with Figure 2.7.2.4).

Figure 2.7.2.6 Daily mean total ozone (bottom solid line) and 11–day running mean total ozone at Halley Station for the 1990s. (The semi–solid line at top is the 11–day running mean for 1957–72 with the dotted lines showing the smoothed extremes during this latter period.)

There is popular speculation that greenhouse warming of two or three degrees over the next century might trigger a similarly large change in association with melting of the two remaining ice–sheets of Greenland and Antarctica. (One scenario of global warming in the context of the “enhanced greenhouse effect” involves a tripling of the concentration of atmospheric carbon dioxide over the next hundred years and, as a consequence, a rise of two or three degrees in the world's average surface temperature. The scenario assumes that the increase of carbon dioxide will level out after a hundred years).

The Greenland ice–sheet is the smaller of the two. It contains an amount of ice which, if melted completely, would raise the level of the world's oceans by 6 m. Calculations suggest a warming of a few degrees might trigger the melting of much of the Greenland ice–sheet. The process would take perhaps one or two thousand years.

The Antarctic ice–sheet is much larger. Its effective volume is equivalent to 55 m of global sea level. It is not expected that it would melt as a result of a warming of two or three degrees. This is because temperatures in most of Antarctica are well below the melting point of ice.

In many places around Antarctica, outward–flowing glaciers fan out into vast ice shelves floating on the ocean. They may disappear quite quickly with global warming of a few degrees. Since they are floating, the process would not directly affect sea level. However it may allow grounded ice behind the shelves to flow faster into the ocean. Calculations suggest this extra flow could contribute one or two metres to sea level over the next one or two thousand years.

In the shorter term – that is, over the next century or two – it is expected that there will be relatively little melting of the ice–sheets. Indeed it is expected that the volume of Antarctic ice will increase slightly because greater snowfall caused by higher evaporation from the warmer oceans will outweigh any increase in melting.

Thus for the next century or two the rise of world–wide sea level will be determined mainly by thermal expansion of the oceans and the melting of non–polar glaciers. (Another possible contributor to sea level change is variation in the volume of water in artesian basins and in the storage of water by people. The uncertainties are large, but it is believed that any contribution to sea level (which might be positive or negative) will be small). The best estimate is a rise at the rate of “several tens of centimetres per century”. (Thermal expansion is likely to continue for at least several centuries until all the world's ocean takes up the increase of surface temperature.) It is difficult to be more precise because, among other things, it is hard to estimate how quickly the warming of the surface will penetrate into the deeper ocean. The relevant processes in the Southern Ocean and in the northern Atlantic are not sufficiently well understood.

Certain regions of the modern world already have experience of rates of sea level change of this order. (Parts of the coastline of Scandinavia are rising about 80 cm per century. Much of the east coast of Japan is falling at 80 or more cm per century. Parts of coastal Holland are falling at about 20 cm per century. Some Pacific islands are rising and some are falling at rates of the order of 10 cm per century, as are parts of the east coast of the United States and Canada. These are regions where the land is rising or falling relative to the sea surface because of geological processes.) "

2.7.3.2 Changes in the Larsen Ice Shelf

Most of the information in Section 2.7.3.2 is reproduced from Hulbe (1997). In her article Hulbe (1997) writes: "There is clearly a connection between warming around the Antarctic Peninsula and the collapse of Peninsula ice shelves. Profound ecological changes are also occurring in response to local climate change. However, temperatures in the interior of the continent have remained fairly constant (Mosely–Thompson, 1992) and it is not yet known whether the observed warming is part of a global trend or is simply a normal fluctuation in local climate. Moreover, warming may actually increase the volume of ice stored in the large Antarctic ice sheets. Dramatic as the retreat of Peninsula ice has been, that ice is less than 1% of the total Antarctic ice volume (Swithinbank, 1988) and its maximum possible contribution to global sea level is less than 50 cm…….……Ice shelves around the northern coasts of the Antarctic Peninsula have been retreating for the last few decades." Hulbe (1997) goes on to point out that an "event to gain attention occurred on the Larsen Ice Shelf, a series of small shelves fringing the eastern coast of the Peninsula from about 71º S to 64º S (Figure 2.7.3.2.1). The seaward front of the northernmost Larsen Ice Shelf (Larsen–A) began a gradual retreat in the late 1940s that ended dramatically in January of 1995, when almost 2000 km² of ice disintegrated into hundreds of small icebergs during a storm (Rott et al., 1996). At the same time, a 70 km × 25 km iceberg broke off the ice front of Larsen–B, between the Jason Peninsula and Robertson Island (Figure 2.7.3.2.2). The settings and styles in which this and other Peninsula ice shelves (for example, the Wordie and Müller Ice Shelves on the western side of the Peninsula (Doake and Vaughn, 1991; Vaughn and Doake, 1996) have retreated are different but the events all coincide with an observed 2.5ºC warming around the Antarctic Peninsula over the last 50 years."

At https://nsidc.org/data/NSIDC-0102/versions/1 it is reported that recent Moderate Resolution Imaging Spectroradiometer (MODIS) satellite imagery analysed at the University of Colorado's National Snow and Ice Data Center revealed that the northern section of the Larsen B ice shelf has shattered and separated from the continent. The shattered ice formed a plume of thousands of icebergs adrift in the Weddell Sea. A total of about 3,250 km2 of shelf area disintegrated in a 35–day period beginning on 31 January 2002.

Hulbe (1997) further explains: "Ice shelves are thick (hundreds of metres) platforms of floating ice that today comprise about 2% of the volume of Antarctic ice. They form where inland glaciers and ice sheets discharge into the ocean. Once afloat, the ice flows by gravity–driven horizontal spreading. Resistance to flow is provided by lateral shear where the shelf flows past bay walls and islands, and by compression upstream of sea–floor rises and islands. The rises and islands are often called "pinning points," a name that belies their importance to the stability of an ice shelf. Ice shelves gain mass by flow from inland ice, by snow accumulation on the upper surface of the shelf, and in some areas by freezing of seawater onto the lower surface. The relative importance of those contributions depends on the size and location of the shelf. Small Peninsula shelves are composed almost entirely of meteoric (snowfall–derived) ice, whereas the bulk of larger shelves, such as the Ross Ice Shelf in West Antarctica, derives from inland ice. Mass is lost primarily by iceberg calving at the seaward ice cliff and secondarily by melting at the lower surface. Except in the northern Peninsula, surface melting makes a trivial contribution to mass loss. Recent estimates of the proportion of ice loss due to calving for all of Antarctica range from 75% to 90% (van der Veen, 1991). The uncertainty is due to the difficulty of measuring melting rates beneath ice shelves. Small iceberg calving (less than 28 km²) is common, whereas larger events occur less frequently. The shelf front may advance for tens of years before an iceberg the size of the 1995 Larsen–B iceberg breaks off. Indeed, the current (1997) Larsen–B front position is similar to its 1960 position on the American Geographical Society map of Antarctica (1981). In the absence of external forcing, the cycle of ice advance and calving maintains a stable shelf front position over time.

Changes in local climate affect ice shelf mass balance (the accounting between the mass of ice gained and the mass of ice lost by the shelf) and thus, the size of the shelf and the location of its calving front. An increase in snowfall, either on the shelf itself or on the inland ice that feeds it, will cause the ice shelf to gain mass. A decrease in snowfall will have the opposite effect. Variations in snow accumulation on the shelf itself are felt quickly, whereas changes in the inland ice may take decades or centuries to be felt. An increase in atmospheric temperature, if large enough to push summer temperatures above the freezing point, will increase mass loss directly by increasing melting at the upper surface. Warmer SSTs that may accompany warmer air temperature could also increase the rate of ice shelf melting. The indirect effects of warmer air temperature are important as well. Melt water, collecting in surface crevasses (wedge–shaped cracks in the ice that normally close by about 50 m depth), allows the cracks to open to greater depth because water exerts a larger outward pressure on crevasse walls than does air. Water–filled crevasses may penetrate to the bottom of the ice, possibly weakening the ice shelf and hastening its decay. Warmer air and sea surface temperatures could produce a positive effect on mass balance by promoting evaporation, which in turn increases snow accumulation over inland ice and the mass of ice flowing into the shelf. Simple comparison between mean summer air temperatures and the locations of extant ice shelves led Mercer (1978) to propose a climatic limit for ice shelf stability: North of the –4ºC mean annual isotherm, ice shelves should be unstable. The present–day warming and retreat of ice shelves around the northern Antarctic Peninsula seem to confirm that hypothesis.

Figure 2.7.3.2.1 A sketch map of Antarctica highlighting the major ice shelves. The heavy line traces the outline of the continent, including grounded ice and bare rock, and the light lines trace the approximate extent of floating ice shelves. (The total volume of ice on Antarctica is 30,109,800 km³, about 2.4% of which is in the form of ice shelves (11% of the surface area). The West Antarctic and East Antarctic Ice Sheets contain 11 and 86% of the total, respectively. Melted, Antarctic ice would contribute 66 m to global sea level (that figure accounts for changes in salinity and for back–filling of the Antarctic continent with water once the ice is removed). Note that melting ice shelves does not increase sea level because the shelves are floating and thus displace a mass of water equivalent to their own mass.)

Hulbe (1997) poses the question: "What does the future hold for Antarctic ice?" and asserts "Unless there is a change in the observed warming trend, further retreats of fringing ice shelves along the Antarctic Peninsula are likely. This seems dramatic on the human scale but is less so on the geologic scale. The present incarnations of Antarctic Peninsula ice shelves have existed for only the last several thousand years and have, in that time, experienced cycles of advance and retreat (Clapperton, 1990; Domack et al., 1995). The present glaciological events may be part of a normal long–term cycle. How ongoing changes to global climate affect the interior of Antarctica remains to be seen. The most sophisticated models available predict an increase, not a decrease, in the volume of ice in the West and East Antarctic ice sheets in future warming scenarios".

Figure 2.7.3.2.2 The northern Larsen Ice Shelf before (left) and after (right) late January

1995. The Antarctic Peninsula and pinning–point islands are shaded (both ice–covered and

bare rock). (The shelf maintains contact with the Seal Nunataks and Robertson Island at present (1997) but new large

shelf–front rifts have been observed (Greenpeace (1997)) and summertime surface melt ponds appear in many satellite

images of Larsen–B. (Scambos and Hulbe (1996)).)

2.7.3.3 Changes in the Ross Ice Shelf

Of more recent interest is the calving of an immense iceberg (B–15) from the Ross Ice Shelf during March 2000. The following information is via the Antarctic Co–operative Research Centre, Hobart, Australia's web home page ( http://www.acecrc.org.au/ ). A June 2000 image of B–15 (and several others) is given in Figure 2.7.3.3.1, while Figure 4.3.3.5.4 shows the location of the iceberg. Figure 4.3.1.2.2. shows the orientation of B–15 in February 2002.

The approximate dimensions of B–15 were assessed in 2000 using AVHRR images as 295 km (~ 170 nm) long, 37 km (~ 25 nm) wide (on average). The area of B–15 was estimated in year 2000 to be about 11,000 km². This is certainly the longest iceberg on record and probably the largest iceberg by area in the modern, satellite era where images from satellites are used to monitor the distribution of icebergs. The previous largest iceberg was A20 (~95 km by 95 km) that calved from the southern Larsen Ice Shelf in early 1986.

In 2000 the thickness of B15 was estimated to range between about 200 m toward the old front and 350 m on the inland margin. Thickness immediately at the old front was probably less than 200 m. Height above sea level of the old front of the Ross Ice Shelf, which now forms one of the long sides of B–15, has been measured to be between 20 m and 40 m. The other long side of B15, which was along the rift in the interior of the shelf, could have been as much as 40 to 60 m in height above the ocean surface.

Figure 2.7.3.3.1 A NOAA–14 infra–red image of B–15, taken at 1341 UTC 9 June 2000. (Courtesy of Matthew Lazzara, Antarctic Meteorological Research Center at the University of Wisconsin – Madison.)

The Ross Ice Shelf front and previous calving events

The front of the Ross Ice Shelf, the seaward margin of the ice shelf, is currently (June 2000) at its most advanced state for any time this past century, and, except for the more easterly section, the most advanced since 1841. A review of this information can be found in Keys et al. (1998). In the sector where B–15 is calving, the front is advancing outwards with the flow and spreading of the ice at a rate of about 900 m/year. So, the 37 km width is equivalent to about 40–45 years of advance of the ice shelf front.

The most recent major calving was of iceberg B–9, 154 km by 35 km, from the sector east of Roosevelt Island, neighbouring the Bay of Whales, and immediately east of the calving site of B–15. B9 calved in 1987, and took 20 months to drift away from the front of the ice shelf, and then across the Ross Sea (Keys et al., 1990).

The cause of the calving of B–15

The calving of B–15 is a natural consequence of the development of an ice shelf. Snow that has fallen on the surface of the ice cover, compacts and forms ice as further snow accumulates on top. The ice gradually flows outwards till it crosses the margins of the grounded ice cover. In many areas this ice flows into ice shelves that are the floating parts of the ice cover. Ice is lost from the ice sheet by calving of icebergs from the outer margins and by melting from the basal surface. The rate of loss from the margins roughly balances the input of snow to the surface.

In satellite images it is possible to see rifts forming in the Ross Ice Shelf many kilometres inland from the margin, and running parallel to the margin. These rifts typically develop and extend over time till an iceberg breaks off. The rifts that form the "calving front" for B–15 could be clearly seen in images acquired in September 1997 over a length of about 240 km. The precursors to these rifts were identifiable many years before this. The calving of B–15 could thus be anticipated, but the exact timing is very difficult if not impossible to predict.

The calving of B–15 is the consequence of a natural progression of events that occur in ice shelves, and is quite unrelated to "global warming" or "greenhouse" effects. Calving events on ice shelves and glacier tongues may occur frequently and produce a few or many small icebergs, or occur rarely and produce one or a few very large icebergs. It is common for many smaller fragments to be produced in a calving event.

2.7.3.4 Disintegration of the Ninnis Glacier tongue

Recently researcher Dr. Rob Massom of the Antarctic Co–operative Research Centre, Hobart has noted the disintegration of the Ninnis Glacier tongue (Massom, 2003): see also https://science.nasa.gov/science-news.