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Polar ice sheets: big, white and cold
A glimpse through the window of a plane during a transatlantic flight shows vast and intriguing, almost mystical, white fields beyond the reach but not the imagination of the common mortal. Vast, they are indeed. The Greenland Ice Sheet is over 2500 km long from South to North and close to 1000 km at its widest, for a total volume of 2.9 106 km3 (Letreguilly et al., 1991). The volume of the Antarctic Ice Sheet is almost ten times as large and the ice-covered region measures over 4000 km across. The gigantic glaciers of Antarctica and Greenland are called “ice sheets” because, with an ice thickness of 2.5–3.5 km and a spread of several thousands of km, their aspect ratio makes them comparable to the sheets of paper used to print this manuscript. Together, the Greenland and Antarctic ice sheets constitute by far the largest present freshwater reservoir and contain enough water to raise global sea level by 70 m if they were to melt, 61 m coming from the Antarctic Ice Sheet (Huybrechts, 2002) and 7 m from the Greenland Ice Sheet (Letreguilly et al., 1991). The large polar ice masses do not merely store water, they play an active role in the climate system.
Actors in the climate system
Ice looks white. A simple fact, but with crucial implications for the radiative budget of our planet, as most of the external energy brought to the Earth comes from the visible radiation of the Sun. Snow and ice have an albedo of 0.6–1, which indicates that they reflect most of the incoming solar radiation; thus, polar regions tend to cool the planet. The phenomenon is amplified in the winter time when large areas are covered in seasonal snow, leading to a positive feedback. The polar energy of drilling sites: DC=Dome C, VO=Vostok, DF=Dome Fuji, BY=Byrd.
Right: Present elevation of the Greenland Ice Sheet simulated with the UBC ice sheet model and ice core sites.
sink also drives the general circulation of the ocean and atmosphere because of the large contrast in temperature between the poles and the tropics, which receive most of the solar energy because of the relative tilt of the Sun’s ray with the normal to the Earth’s surface: energy absorbed in the tropics is exported to higher latitudes through atmospheric and oceanic flow. The positive feedback on the radiative budget is also excited during glacial periods, when large ice sheets were stimulated to grow and cover extensive areas by a decrease in incoming solar radiations above the Arctic Circle, as explained by the Milankovitch theory.
Astronomic theory of climate change
The relative position and orientation of the Earth during its yearly elliptical journey around the Sun determines how much solar radiation is received at different latitudes during a year. The received radiation is controlled by three orbital parameters described in Fig. 2.3: the obliquity or tilt of the Earth polar axis relative to the elliptic plane, the precession of equinox that corresponds to the season when the Earth is closest to the Sun on the ellipse (perihelion), and the eccentricity that describes the degree of elongation of the ellipse. Tilt causes seasons because the hemisphere pointing toward the Sun receives more radiation. Change in tilt between 22 and 25 occurs with a 41 kyr (thousand years) periodicity and controls the intensity of seasonality: the larger the tilt, the larger the insolation contrast between the two hemispheres. Today the Earth reaches the perihelion in January, causing slightly milder winters in the Northern Hemisphere, as opposed to 11 kyr ago when perihelion was approached in July (19 kyr and 23 kyr periodicity). Eccentricity varies in 100 kyr and 400 kyr cycles and controls how the timing of perihelion combines with seasonality (tilt); thus, the relative severity of summer and winter time in each Hemisphere.
Milankovitch (1930) calculated the past position and inclination of the Earth and associated the changes in seasonality with the glacial–interglacial cycles that have characterized the Earth’s climate for the past 2 Myr. Ice ages, cold temperatures and large ice volume, correlate well with the 65 N summer insolation (Emiliani, 1955; Berger, 1978), as shown on the Vostok record Fig. 2.2; thus, small changes in insolation get amplified by feedbacks in the climate system that lead to a global cooling of 5 C and the inception and growth of large ice sheets that drew down sea level by 120– 135 m (e.g., Yokoyama et al., 2000). If insolation is indeed the “pacemaker of ice ages” (Hays et al., 1976), future glaciations can be predicted and the next one should not happen before at least 50 kyr (Berger and Loutre, 2002). However, the astronomical theory of climate does not tell how warm the climate is going to be nor how high sea level will rise over the next decades and centuries; thus, I will especially focus on the climate and behaviour of ice sheets during former periods that most resembled modern conditions and try to improve the understanding of glacial archives by predicting the conditions under which they were generated and preserved.
Ice sheets and fast climate change
Though the Milankovitch theory may well explain the inception and expansion of large ice sheets, isotopic and geomorphologic records display a high frequency variability during glacial periods that is not present in the insolation signal. For instance, deep-sea sediment records in the North Atlantic (Fig. 2.4) contain a series of layers that are rich in sediments also found in eastern Canada and poor in foraminifera, thus reflecting cold sea-surface temperature. These records suggest that intense and short-lived release of large armadas of icebergs released from the Laurentide Ice Sheet exported enough freshwater to shut down the thermohaline circulation of the ocean, effectively cooling down the climate in the North Atlantic region by 5 C. These Heinrich Events (Heinrich, 1988; Bond et al., 1992) are believed to result from surges or oscillations in the ice stream regime of the Laurentide Ice Sheet and have motivated the development of the UBC Ice Sheet Model used for my studies of Greenland (Marshall, 1996). These events reflect the dynamic role of internal ice dynamics on the global climate and stimulate further research to better understand the flow of ice sheets.
A brief history of ice core drilling
The potential of the glaciological archive was understood about 60 years ago, but the remoteness and harsh climatic conditions of polar ice sheets and the resulting technical and financial burden have so far limited the use of the “glaciological mine” (Reeh et al., 1987) to a dozen ice cores. Scientific expeditions aimed at understanding the climate in Antarctica started in 1957–1958 during the International Geophysical Year, followed the next year by a 1400-km traverse organized by the Americans and a decade later by deep drilling operations.
The first deep ice core campaign was inaugurated in Northwest Greenland at Camp Century. Bedrock was reached in 1966, with a 1390-m deep core that spans 100 kyr of climatic history (Dansgaard et al., 1969). In Antarctica, 2000 m of ice containing 75 kyr of history were extracted at Byrd near the Ross and Amundsen Sea in 1968 (Epstein et al., 1970). In the early 1970’s, the Russians installed a station at Vostok in central east Antarctica, one of the coldest places on our planet with an average temperature of 55 C and extreme temperature of 89 C. Several ice cores were drilled and successively reached 500 m in 1970, 950 m in 1974 (Barkov and Gordienko, 1976) and much deeper more recently. Also in East Antarctica, the French extracted a 900-m core at Dome Concordia in 1978 (Lorius et al., 1979). Drilling returned to Greenland and in 1981, the Danes, Swiss and Americans pulled a 2000-m ice core at Dye 3 in the southern part of the ice sheet. The record contains over 100 kyr of climatic history (Dansgaard et al., 1982), but the lower part of the record is difficult to interpret because it is strongly affected by ice flow, an inescapable problem for flank flow sites, an issue I will try to address with three-dimensional modelling of ice flow.
Drilling resumed at Vostok in 1982 with a Russian–French association that benefited from American logistic support. High expectations were put on this operation because the low surface ac-cumulation rate, only 2 cm per yr, raising the possibility that the age of the ice might greatly exceed that for the previously drilled sites. Efforts were soon rewarded with a 2000-m ice core spanning 150 kyr of climate history (Lorius et al., 1985), providing the first detailed record reaching beyond the last glacial cycle and into the next one, embracing an interglacial period somewhat similar to modern conditions about 125 kyr ago, an epoch called Eemian or Marine Isotope Stage 5 (MIS 5) in reference to the chronology of marine isotopic records (Martinson et al., 1987). Drilling reached a final depth of 3623 m in 1998, just 80 m above the large subglacial Lake Vostok, with 3310 m of ice spanning 420 kyr of climatic history (Petit et al., 1999, and Fig.2.2), the bottom 300 m originating from basal freeze-on of water from Lake Vostok. The record thus contains four glacial–interglacial cycles and reaches MIS 11 (395–420 kyr BP), a long interglacial period of special interest because its astronomical parameters resemble modern conditions (Drowler et al., 2002).
In the early 1990’s, Europeans and Americans drilled two deep ice cores 28 km away from each other near the 3300-m-high Summit of the Greenland Ice Sheet. The European GRIP project reached the 3029-m-deep bedrock in July 1992 (GRIP Members, 1993), the American GISP2 project 3087 m the following summer (Grootes et al., 1993). Their deep ice from the last inter-glacial period led to much speculation about rapid climate variability. However, comparison of greenhouse gases with the Vostok record clearly proved the Greenland cores had been corrupted by strong flow disturbances that had mixed ice deposited during the Eemian (Chappellaz et al., 1997). A new drilling program aiming at better sampling Eemian ice in Greenland was started in 1996 at NorthGRIP, 300 km North of the Summit sites. Bedrock was finally reached in 2003 and appears to contain ice up to 123 kyr BP, that is ice reaching into a stable part of the Eemian (NGRIP Members, 2004).
Besides the Vostok program, other deep ice cores have recently been drilled in Antarctica. The Japanese went to Dome Fuji, 1500 km away from Vostok and closer to the Atlantic Ocean, and extracted 2503 m of ice before the drill got trapped into the ice. The ice core record showed three glacial cycles strikingly similar to Vostok, providing evidence of an homogeneous climate over most of Antarctica (Watanabe et al., 2003). Europeans started a new deep drilling project (EPICA) at the French–Italian base of Dome C (DC) in 1998. In 2003, they recovered 3190 m of ice spanning the longest-to-date ice-core record with up to eight glacial cycles of climatic history (EPICA community members, 2004).
Methods for dating ice
Although annual ice layers can preserve their original chemical content, the climatic history can be obscured by the thinning and deformation processes resulting from ice flow; thus, determining the age of ice can be a real challenge. Annual layers tend to merge at greater depth, especially in Antarctica where initial layers are already thin because of low surface accumulation. As previously seen, flow disturbances have also particularly affected sites like GRIP and GISP2. Each ice core record has its own flow and accumulation rate characteristics, therefore many strategies have been applied to recover the essential chronology of events.
I distinguish three main classes of methods for extracting the age of ice along a core: (1) Count-ing of annual layers, which is accurate (Alley et al., 1997) if layers are thick enough. It has been applied for dating the youngest 37.8 kyr of the GISP2 record (Meese et al., 1997) and from 8.2 to 11.5 kyr BP for GRIP (Johnsen et al., 1997). Accumulation rate is too low in central Antarctica for this method. (2) Use some level of comparison to a chronology or to dated events, either by referring to time markers (e.g., 10Be peak, 14C, or U-Th Genty et al., 2003) and other dated time series, or by applying an orbital tuning by linking the measured changes in the core to the well known insolation signal (Berger, 1978). Time markers are useful but limited to 10’s kyr. Orbital tuning assumes a linear response of the climate system to change in insolation, introducing a 6 kyr uncertainty on age (Parrenin, 2002). The method has been widely used both on ice and marine core records (e.g., Meese et al., 1997; Martinson et al., 1987; Bassinot et al., 1994). Chronologies can also be syn-chronized by correlating the isotopes of atmospheric oxygen, a method also compatible with marine records because the composition of the ocean and the atmosphere change simultaneously (Bender et al., 1994), or by methane concentration to compare Antarctic and Greenland records (e.g., Blunier and Brook, 2001), because methane is well mixed in the atmosphere. (3) Ice flow modelling, which requires assumptions concerning the simplicity of ice flow, on past surface elevation and origin of ice, and an accumulation model. This method is satisfying because it includes the physics of ice, but is highly dependent on the accumulation scenario, which is generally tied to the isotopic variations in the record. Such relations have been shown to fail at particular periods (Cuffey and Clow, 1997), explaining for instance the continuing debate on the chronology of Summit ice cores (Grootes et al., 1993; Meese et al., 1997; Johnsen et al., 1995, 2001; Shackleton et al., 2004). Vostok experiences all the difficulties of the modelling approach because of a distant origin of deep ice (Ritz, 1992), rough underlying topography (Siegert and Kwok, 2000) and possible hiatus in surface accumulation (Parrenin et al., 2001). The methods can be combined (e.g., GRIP-SS09 chronology, Johnsen et al., 2001) and Parrenin (2002) has recently developed a highly promising “federative approach” that picks up the best features of all these dating techniques to obtain the age–depth relationship of the Vostok, DC and Dome Fuji records.
Table of contents :
CHAPTER 1: Introduction
1.1 Context and objectives
1.2 Thesis outline
CHAPTER 2: Polar ice sheets: actors in the climate system, archives of the past 5
2.1 Ice sheets and climate
2.1.1 Polar ice sheets: big, white and cold
2.1.2 Actors in the climate system
2.1.3 Astronomic theory of climate change
2.1.4 Ice sheets and fast climate change
2.2 Ice core records
2.2.1 A brief history of ice core drilling
2.2.2 Methods for dating ice
2.2.3 What do we learn from ice cores?
2.2.4 Past and future climate
2.3 Inferring past temperature from water isotopes
2.3.1 Stable water isotopes
2.3.2 Isotopic fractionation
2.3.3 Fractionation during precipitation
2.3.4 Fractionation during evaporation
2.3.5 Rayleigh distillation
2.3.6 Principle of isotopic thermometry
2.4 Conservation of the water-isotope signal in snow and ice
2.4.1 Diffusion in the firn
2.4.2 Diffusion in ice
2.5 Modelling ice-sheet evolution
2.5.1 Dynamics of ice flow
2.5.2 Principles of ice sheet modelling
2.5.3 Parameterization of the climate over ice sheets
CHAPTER 3: Three-dimensional tracer modelling
3.1 Motivation
3.2 Transport of provenance variables
3.2.1 Technical context
3.2.2 Tracer principles
3.2.3 Eulerian, Lagrangian, or semi-Lagrangian scheme?
3.3 Specificity of the age–depth relationship
3.3.1 Sources of error
3.3.2 A balance-based interpolation method
3.3.3 Age boundary conditions
3.3.4 Three-dimensional implementation
3.4 Depositional model
3.4.1 An archive like a “birth registry”
3.4.2 Present isotopic distribution
3.4.3 Past isotopic distribution
3.5 Practical application
3.5.1 Tracer transport flow chart
3.5.2 Construction of tracer stratigraphy
3.5.3 Global properties
3.5.4 Average isotopic concentration
3.5.5 Application of tracer modelling
CHAPTER 4: Global stratigraphy of the Greenland Ice Sheet
4.1 Introduction
4.2 Model and experiment
4.2.1 Ice dynamics
4.2.2 Climate forcing
4.2.3 Practical details
4.3 Depositional provenance stratigraphy
4.4 Stratigraphy at ice core sites
4.4.1 Isotopic stratigraphy
4.4.2 Borehole catchment
4.5 Age of Greenland ice
4.5.1 Age of deep ice
4.5.2 Average age of the ice sheet
4.6 Conclusion
CHAPTER 5: Constraints on Greenland ice cores and glacial history
5.1 Introduction
5.2 Validation process and effect of model parameters
5.2.1 Methodology and validation process
5.2.2 Best parameters
5.2.3 Sensitivity of parameters
5.3 Modelling the GRIP records
5.3.1 Age–depth profile
5.3.2 Borehole temperature
5.3.3 Paleotemperature
5.3.4 Paleo-elevation
5.3.5 Ice origin and ice-divide migration
5.3.6 Constraints from GRIP
5.4 Constraints from other Greenland cores
5.4.1 GISP2
5.4.2 Dye 3
5.4.3 Camp Century
5.4.4 NorthGRIP
5.5 Minimal configuration during the Eemian
5.6 Conclusion
CHAPTER 6: Global stratigraphy of the Antarctic Ice Sheet
6.1 Description of the model
6.1.1 Ice dynamics
6.1.2 Climate forcing
6.2 Experimental design
6.2.1 Adjustments near ice core sites
6.2.2 Model spin-up
6.2.3 Validation process
6.3 Depositional provenance stratigraphy
6.3.1 “East–west” profiles
6.3.2 “North–south” profiles
6.4 Deposition age
6.4.1 Age stratigraphy
6.4.2 Age of deep ice
6.4.3 Average age of the ice sheet
6.5 Isotopic stratigraphy at ice core sites
6.5.1 Dome C
6.5.2 Vostok
6.5.3 Dome Fuji
6.6 Conclusion
CHAPTER 7: Simple vs. complex ice flow models for deep Antarctic records
7.1 Introduction
7.2 Prediction of age–depth in an ice core
7.2.1 Ice flow dating model
7.2.2 Age–depth for the ice sheet model
7.3 Depositional conditions for the deep East Antarctic records
7.3.1 Surface mass balance
7.3.2 Ice origin
7.3.3 Paleo-elevation of depositional ice
7.4 Internal ice flow properties at deep ice core sites
7.4.1 Thinning
7.4.2 Velocity profiles at domes
7.4.3 Velocity profiles at Vostok
7.4.4 Simple vs. 3D calculation of velocity
7.5 Discussion and conclusion
CHAPTER 8: Global budget of water isotopes inferred from polar ice sheets
8.1 Introduction
8.2 Conditions for an accurate diagnosis
8.2.1 Ice sheet reconstruction
8.2.2 Deposition rate of water isotopes
8.2.3 Stability of the Antarctic Ice Sheets
8.2.4 Presentation mode
8.3 Volume and composition of the East Antarctic Ice Sheet
8.3.1 Ice volume and sea level
8.3.2 Isotopic composition of the EAIS
8.3.3 Effect of the EAIS on ocean composition
8.4 Volume and composition of the West Antarctic Ice Sheet
8.4.1 Ice volume and sea level
8.4.2 Isotopic composition of the WAIS
8.4.3 Effect of the WAIS on sea water composition
8.5 Volume and composition of the Greenland Ice Sheet
8.5.1 Ice volume and sea level
8.5.2 Isotopic composition of the GIS
8.5.3 Effect of the GIS on sea water composition
8.6 Discussion: ice sheets and sea water composition
8.6.1 Sea level vs. sea water composition
8.6.2 Last Glacial Maximum
8.6.3 In an ice-free world
8.6.4 Interglacials
8.7 Conclusion
CHAPTER 9: Conclusion
9.1 Tracers
9.2 Glaciological results
9.3 Suggestions for future work
REFERENCES .
