- Email: [email protected]
ICE CORE RECORDS | Greenland Stable Isotopes
Greenland Stable Isotopes B M Vinther and S J Johnsen, University of Copenhagen, Copenhagen, Denmark ã 2013 Elsevier B.V. All rights reserved.
The Early Isotope Work By doing stable isotope investigations of frontal precipitation over Copenhagen, Denmark, in the early 1950s, Willi Dansgaard discovered a correlation between condensation temperatures and isotopic ratios. Subsequently, Dansgaard proposed in a groundbreaking paper that paleoclimate could be studied by doing stable isotope measurements on Greenland ice cores (Dansgaard, 1954). At this time, no ice cores were available from Greenland for such investigations, and the immense scientific scope of this proposal was hardly realized even 15 years later when the first results from the Camp Century core became available. Carl Benson of the US Army Snow, Ice, and Permafrost Research Establishment (SIPRE) was, however, already doing pit studies in Northwest and Central Greenland in the early and mid-1950s (Benson, 1962). In cooperation with Epstein and Sharp, he discovered the very strong annual isotopic cycle in the Greenland snow pack. Eventually, this cycle became most useful for measuring accumulation rates and dating ice cores. When the Site 2 ice core (see Figure 1), which was drilled by SIPRE in preparation for the International Geophysical Year’s Antarctic activities in 1957, became available in 1955, Chester C Langway, also from SIPRE, initiated the modern era of multidisciplinary ice-core studies including detailed isotope profiles deep down in the core (Langway, 1967). The Site 2 isotope data (Epstein and Sharp, 1959) showed a clear annual cycle below the firn layers, but smoothed compared with the isotope data from Benson’s pit studies (Langway, 1967). Thus, Langway discovered the effect of firn diffusion discussed below. Willi Dansgaard continued his studies in Greenland and discovered the isotopic latitude effect for Greenland and the strong correlation between mean annual temperatures and isotopic values on the ice cap (Dansgaard, 1961). He also used the Rayleigh condensation model to improve the understanding of the physical basis for the climatic information contained in natural precipitation. Most of his earlier results are summarized in his classic Tellus paper (Dansgaard, 1964) where he also provided the first results from the International Atomic Energy Agency (IAEA) precipitation network and introduced the concept of deuterium excess. The Camp Century deep ice core was recovered by Lyle Hansen and his coworkers from the US Army Cold Region Research and Engineering Laboratories in 1966. Two years later, Dansgaard suggested that Langway do a comprehensive isotope study on the core, a proposal that Langway accepted. This most fruitful cooperation was later extended with the group of Hans Oeschger from Bern, Switzerland, and became the basis for the Greenland Ice Sheet Program (GISP) that managed to recover a deep core from Dye-3 South Greenland with a new Danish deep drill in 1981. When the first Dye-3 isotopic results were published (Dansgaard et al., 1982), the immense impact of the isotopic paleostudies on Greenland
deep ice cores was firmly established. In the early 1970s, the Danish group also built a unique inlet system for running 512 daily mass-spectrometric oxygen-isotope measurements. The original aim was to date the Dye-3 core by counting annual isotopic cycles through most of the Holocene by measuring a total of 80 000 samples. However, this system greatly strengthened the isotope work done with subsequent Greenland ice cores and over one million samples were analyzed using it.
Shallow Ice Cores – The Role of Diffusion A large number of ice cores reaching depths of 100–200 m have been drilled at various locations on the Greenland ice sheet. Such cores offer a time span ranging from a few centuries to thousands of years, depending on core depth and local accumulation rates. As instrumental climatic observations are limited to the past two centuries for most of the Northern Hemisphere, the records of d18O from shallow ice cores are a valuable asset when investigating relationships between climate and d18O. Due to the multiple cores available, it is possible to stack the records and thereby reduce the glaciological noise in the d18O data. Using cores from various locations, it is also possible to resolve local climatic variations affecting only parts of the ice sheet. The investigation of d18O records from the upper part of the ice cores is complicated by water vapor diffusion in the pores of snow and firn in the top 60–70 m of the ice sheet. This diffusion tends to smooth the high-frequency oscillations of the d18O. Figures 2 and 3 show the diffusion-related decay of the annual d18O cycle for a central Greenland shallow core.
The Climate Signal in Seasonal d18O The close relationship between d18O and site temperature makes d18O a valuable proxy both for Greenland climate and Northern Hemisphere atmospheric circulation patterns. To exploit the potential of Greenland d18O records fully Vinther et al. (2003, 2010) have shown that it is essential to consider seasonal oscillations in the d18O records. In the vast areas of the Greenland ice sheet where annual accumulation exceeds 20 cm of ice equivalent, seasonal oscillations of d18O can be retrieved through most of the Holocene. In areas of lower accumulation, the seasonal d18O cycles are obliterated by the diffusion in the upper snow and firn layers. As can be seen in Figures 2 and 3, the seasonal cycle of d18O is influenced by diffusion even in areas with accumulation rates larger than 20-cm ice equivalent per year. In the top 60–70 m of the ice sheet, a marked decay in amplitude of the d18O cycle can be observed. If a climatic interpretation of seasonal d18O values is to be pursued, it is necessary to
403
ICE CORE RECORDS | Greenland Stable Isotopes
understand the diffusion process that masks the seasonal oscillations in the d18O signal.
−90
−80 −70 −60 −50 −40 −30 −20 −10
0
10
80
80
76
76
72
72
68
68
64
64
60
60 −50
−40
−30
d 18O (per mil)
Figure 1 Locations of selected Greenland ice cores. The Site 2 core was the first ice core drilled in Greenland. The Camp Century, Dye-3, GISP2, GRIP, NGRIP, and Renland cores all contain ice from the Holocene, Glacial, and Eemian periods.
In a publication by Johnsen (1977), the first step was taken in comprehending the physics behind the diffusion process. Later papers by Merlivat and Jouzel (1979), Herron and Langway (1980), and Schwander et al. (1988) led to the Johnsen et al. (2000) paper in which a physical model covering both d18O and dD diffusion was presented and validated. The Johnsen et al. (2000) model predictions of the diffusion-driven dampening of the annual d18O cycle can be seen in Figure 3. Using the Johnsen et al. (2000) diffusion model and the mathematical formulation of the diffusion problem presented in Johnsen (1977), it is possible to correct d18O series for diffusion. An example of diffusion-corrected d18O data can be seen in Figure 2(b). Having diffusion-corrected data, the underlying seasonally resolved climatic signal can be retrieved from ice-core d18O records. It is easy to demonstrate the importance of correcting the seasonal d18O records for diffusion before any climatic interpretation is pursued. Four different central Greenland ice cores were drilled in 1984/1985 about 500 km from the coastal Tasiilaq meteorological station. Table 1 demonstrates that correlations between winter season d18O and December-to-March temperatures at the Tasiilaq station (Laursen, 2003) are greatly improved by diffusion correcting the records. Furthermore, it can be seen that the ice-core records from the areas of modest accumulation (Sites E and G) benefit the most from the diffusion correction. Using the approach of diffusion correction on several Greenland ice cores, Vinther et al. (2003, 2010) demonstrated that winter season d18O from Greenland ice cores contained both a distinct Greenland temperature signal and the signature of the North Atlantic Oscillation (NAO), a large-scale atmospheric flow pattern. In fact, Greenland annual average temperatures were found to be better captured by winter season d18O data than annual average d18O data. A linear regression between the multi-ice core winter d18O signal and Northern Hemisphere December-to-March sea level pressure (slp) is seen in Figure 4. The NAO pattern with opposite pressure trends over Iceland and the Azores is clearly seen in Figure 4. The regression was carried out for the years
−28
−28
−30
−30
−32
−32
−34
−34
−36
−36
−38
−38
−40
2
3
4 5 Depth (m)
(a) 18
6
7
70 (b)
70.5
71 71.5 Depth (m)
72
d 18O (per mil)
404
−40 72.5
Figure 2 Measurements of d O from different depths in the Site E central Greenland ice core. Left: near the surface before firn diffusion dampens the annual cycle. Right: below the diffusion zone, the annual cycle is significantly dampened. The thin curve is diffusion-corrected data. Annual accumulation at Site E amounts to 21 cm of ice equivalent.
ICE CORE RECORDS | Greenland Stable Isotopes
405
10 20
Depth (m)
30 40 50 60 70 Observed Modeled
80
0
0.5
1
1.5 2 2.5 Amplitude (per mil)
3
3.5
4
Figure 3 Amplitude of the annual d18O cycle in the Site E central Greenland ice core and model estimate.
Table 1 Correlations (r) of winter d18O with Tasiilaq December-toMarch temperatures (1895–1982) Ice core
Site A
Site B
Site E
Site G
Mean ABEG
Accumulation (ice equivalent; cm) r (before diffusion correction) r (after diffusion correction)
29
31
21
25
–
0.29
0.43
0.25
0.21
0.36
0.33
0.50
0.41
0.37
0.52
1872–1970 and the slp-data are from the HadSLP1 dataset (Basnett and Parker, 1997).
Deep Ice Cores – Climate During 125 Millennia During the past 40 years a total of five deep ice cores have been drilled all the way through the Greenland ice sheet. All five records contain ice from the entire Holocene, the last glaciation, and the Eemian. A sixth ice core drilled through the small east coast Renland ice cap also reaches back into the Eemian (for drill site locations see Figure 1). Measurements of d18O have been carried out on the full length of all six ice-core records. From Figure 5 it is clear that the six d18O records are very similar for the past 40 000 years. Below 40 000 years, the Dye-3 record is disturbed by ice flow, as is the Renland core below 60 000 years. The Greenland Icecore Project (GRIP) and GISP2 records are affected by flow-induced foldings below 105 000 years. The North Greenland Icecore Project (NGRIP) core presents the only d18O record known to be stratigraphically continuous into the Eemian. However, Eemian ice is found in all six ice cores.
Timescales for Deep Ice Cores A prerequisite for the climatic interpretation of the deep ice cores is having a well-established depth–age relation. For most
Figure 4 December–March slp-response to a one-sigma increase in winter d18O in a series based on multiple Greenland ice cores. Contour interval is 0.4 hPa, red contours are positive while blue are negative.
of the Holocene, counting annual layers using the seasonal oscillations in d18O is a very robust way of obtaining a timescale for the Greenland ice cores. During the glacial, generally lower accumulation rates lead to diffusion-caused obliteration of the annual d18O cycles, rendering annual layer counting impossible. The first approach to date the glacial part of the ice cores relied on analytical 2D flow model estimates of the progressive thinning of annual layers (Dansgaard and Johnsen, 1969). The flow model presented by Dansgaard and Johnsen (1969) (D–J model) proved to be very adaptive and Johnsen and Dansgaard (1992) refined the analytical D–J model to account for both increased shear near the bottom of the ice sheet and sliding of the ice sheet over the bedrock. The D–J model can also be modified to include flow due to bottom melting affecting the ice sheet. To include temporal variations of the accumulation rate, a numerical version of the D–J model was required. The first effort to create such a model was made by Johnsen et al. (1992). Numerical D–J flow models have been used to provide timescales to both the GRIP and the NGRIP ice cores (NGRIP members, 2004). These models require past accumulation rates as an input parameter. This requirement can be fulfilled by using the relationships between d18O and accumulation rate at the drill sites. Figure 6 shows the Dye-3, GRIP, and NGRIP d18O profiles on timescales derived from numerical D–J flow models. As the bottom-most part of the Dye-3 core is disturbed by ice flow, only the past 40 000 years of the Dye-3 profile is shown. The similarity of the three profiles is striking. The first effort to count annual layers during the glacial was carried out on the GISP2 core (Meese et al., 1997). Seasonally changing features in the visual stratigraphy of the GISP2 core were used for counting the layers. Surprisingly, the glacial GISP2 timescale derived from this methodology did not show a persistent relationship between the d18O and the
406
ICE CORE RECORDS | Greenland Stable Isotopes
Greenland deep ice core isotopes Camp century
Dye-3
GISP2
GRIP
Renland
NGRIP
305
10
1600 1600
1150
1
1800
1500
1
1
1
1
1800
15
1
1800 2 4
Depth (m)
14
2200
8
4
4
4
8
2200
1900
2400
19
12
14
12 14
315
50 14 2500
19
2600
2600
21
23
23
2800
2800
3000
3000
2000
21
320
3000
−35 −30 −25
−45 −40 −35 −30
−45 −40 −35 −30
60
80
21 23
1350
−45 −40 −35 −30 −25
18 20
21
23
40
12
14
2400
18
1950
1300
8
14
8
30
2000
12
8
12
21
2
20
NGRIP age (ka)
4
12
2
2000
4
8
1250
2
2000
2
1850
1200
2
310
−30 −25
23 24 25
100
120
−45 −40 −35 −30
d18O (‰)
Figure 5 Early Holocene, Glacial, and Eemian d18O records from six Greenland ice cores are shown. The cores are all on depth scales. The NGRIP timescale is shown, and the numbers on the left side of the scale give the most prominent interstadials.
Dye-3, GRIP and NorthGRIP isotopic profiles 1
−30
2
4
GRIP NorthGRIP Dye-3
8
d18O (‰)
Eem
−35
12
14
18
20
21
23
24
25
−40 −45 0
20
40
60
80
100
120 × 103
Age (GRIP2001 years b2k)
Figure 6 Profiles of d18O from the Dye-3, GRIP, and NGRIP cores, on the ss09sea flow-modeled timescale. Numbers are given for the most prominent interstadials. b2k, before AD 2000.
accumulation rate back in time. Such a relation is underpinned both by modern observations and the following simple argument: a change in d18O is associated with a change in site air temperature, and as the air temperature determines the maximum water vapor content of the air, the amount of moisture reaching the site will be related to observed d18O values. Based on a new effort to date the glacial part of the NGRIP core using both seasonal oscillations in chemical impurities and visual stratigraphy, it is now believed that the apparent breakdown of the d18O/accumulation relationship observed in the GISP2 core is an artifact caused by erroneous
interpretation of the GISP2 visible strata (Andersen et al., 2006; Svensson et al., 2006).
The Holocene The major part of the deep ice cores consists of Holocene ice. Looking at Figure 7, it is obvious that millennial scale trends in the ice core records are different during the course of the Holocene. Vinther et al. (2009) have shown that these differences are most likely due to a combination of past Greenland ice sheet elevation change and upstream effects (i.e., that ice in
ICE CORE RECORDS | Greenland Stable Isotopes
407
2
d18O deviation (per mil)
0
-2
Camp century NGRIP GRIP GISP2 DYE-3
-4
Agassiz/Renland
-6
0
2000
4000
6000 Age b2k (years)
8000
10 000
Figure 7 Holocene profiles of d18O deviations with respect to present. Heavy lines show millennial scale trends in the data. All profiles are shown on the GICC05 time scale, except for GISP2 d18O which is shown on the original GISP2 time scale. b2k, before AD 2000 .
the lower strata originates from higher elevation positions upstream from the drill site). Furthermore, Vinther et al. (2009) demonstrated, that the average d18O from the two small ice caps of Renland (eastern Greenland) and Agassiz (Ellesmere Island north west of Greenland) can be used to establish the Holocene climatic d18O evolution, if corrected for post-glacial rebound (the corrected average d18O is shown as the black curve in Figure 7). The differences between the Renland/Agassiz d18O average and d18O from the other five ice cores shown in Figure 7 can thus be converted directly to past changes in depositional elevation at the five locations using the observed d18O-elevation relationship for Greenland. A conversion of the Renland/Agassiz d18O average to past Greenland temperatures show a climatic optimum 2–3 C warmer than present (Vinther et al., 2009) in good agreement with results from Dahl-Jensen et al. (1998) showing that borehole temperature measurements at GRIP and Dye-3 imply that temperatures on the ice sheet were, 3 C higher than present during the climatic optimum. All but one of the d18O records shown in Figure 7 are plotted on the Greenland Ice Core Chronology 2005 (GICC05). GICC05 is a newly counted timescale that synchronizes the ice cores using common volcanic reference horizons (Rasmussen et al., 2006; Vinther et al., 2006, 2008). Exploiting the synchronization of the d18O records (Rasmussen et al., 2007) showed that the only significant common events in the early Holocene Dye-3, GRIP, and NGRIP records are the decreases in d18O some 8200, 9300, and 11 400 years ago. These events are known as the 8.2 ka and 9.3 ka cold events and the Preboreal oscillation, and they are observed in ocean sediment cores, tree ring data, and stalagmite records from all over the Northern Hemisphere. The events are believed to be associated with very large freshwater outbursts from the decaying Laurentide ice sheet. The large amount of freshwaters would temporarily weaken the North Atlantic overturning
circulation thereby diminishing the oceanic heat transport towards the polar region.
The Last Glaciation The abrupt shifts in d18O leading to the warm Greenland interstadials (Dansgaard et al., 1993; Johnsen et al., 1992) are often referred to as Dansgaard–Oeschger events. These events are the most prominent features observed in the glacial parts of all Greenland ice cores (see Figures 5 and 6). The shifts are most probably due to large-scale reorganizations of the atmospheric and oceanic circulations during the glacial. Hence, full glacial (stadial) conditions are probably related to a shutdown of the North Atlantic overturning circulation, while the interstadials are warmed by the return of the overturning circulation. Changes from stadial climatic conditions to the 10–15 C warmer interstadial climate are believed to have happened within a decade or two. The glacial climate is therefore characterized by a degree of instability that cannot be found during the Holocene. This instability is likely to be caused by the N–S shifts of the atmospheric polar front that follows the oceanic polar front (Ruddiman and McIntyre, 1981) in the glacial. The d18O at the polar front is close to 8% Dansgaard and Tauber (1969), which is also likely valid in the glacial. The Greenland d18O values are mainly driven by the latitudinal gradient of 0.54% d18O/lat, the distance to the polar front and elevation changes Dansgaard (1961) and Johnsen et al. (1989). Johnsen et al. (1997) estimated the d18O short-term annual RMS (Root Mean Square) variability from the GRIP ice core (Figure 5 in that paper). The records of annual variability (Figure 5(c)) and the d18O (Figure 5(a)) are in anticorrelation. The RMS variability for Holocene, interstadials, and stadials are 0.9%, 1.8%, and 3.4%, respectively. The increased variability is partly due to lower accumulation rate, in the interstadials and stadials, and partly due to the rapid N–S shifts of the polar front during these cold periods.
408
Table 2
ICE CORE RECORDS | Greenland Stable Isotopes Average values and differences of d18O during different climatic periods (all values are given in %)
Drill site Camp century Dye-3 GISP2 GRIP NGRIP Renland
Present (P) 29.5 27.6 35.0 35.1 35.4 27.3
Stadial (S) 41.9 36.1 42.2 41l.9 44.5 32.1
Johnsen et al. (1992) found that an intercomparison between glacial d18O values observed in different Greenland ice cores shows important differences in the size of the d18O increase associated with the shift from stadial to Holocene climatic conditions. The Camp Century core shows the largest d18O changes while the changes at Renland are the most modest (see Table 2 and Figure 5). It is most likely that the large changes in Camp Century (and NGRIP) d18O are related to elevation changes, but the changes could also be partly due to a strongly depleted northwesterly source of precipitation during the stadials at those sites. Such a source would stem from a split in the polar jet-stream, with branches both north and south of the Laurentide ice sheet. Precipitation arriving from the northwest would not be able to reach the Renland ice cap as it is situated east of the main ice sheet, and therefore, is shielded from any such precipitation. Hence, the relatively modest decrease in Renland d18O values during the stadials can be understood in terms of the Renland ice cap not receiving precipitation from the depleted northern branch of a split jetstream. Recent atmospheric modeling efforts suggesting a split jetstream during the glacial in the winter (i.e., Bromwich et al., 2004) therefore enjoy significant support from observations of stadial d18O in Greenland ice cores. The whole glacial period in Camp Century is 3% colder compared to the present than seen in NGRIP. This difference is most likely caused by 500– 600 m higher elevation at the Camp Century compared to the NGRIP elevation.
The Eemian Ice from the last interglacial, the Eemian, has been found near the bedrock in all deep ice cores retrieved from the Greenland ice sheet (see Figure 5). It is, therefore, safe to conclude that most of Greenland was covered by ice during the Eemian. By using the d18O values observed in Eemian ice, it is possible to further quantify the changes in the size of the Greenland ice sheet during the Eemian. Comparing present day d18O values to those observed in the Eemian ice, it is clear that the Eemian had 2.5–3.5% higher d18O values than presently observed all across Greenland (Table 2 and Figure 5) except for at the southern Dye-3 site. At Dye-3, Eemian d18O was 4.7% higher than recent values. The observed d18O value at a site on the Greenland ice sheet is dependent on elevation (the gradient is approximately 0.6% per 100 m altitude according to Johnsen et al., 1989). It can therefore be inferred from the rather uniform (3%) increase
Eemian (E) 27.0 22.9 32.3 31.6 32.3 23.8
P–S
E–P
12.4 8.5 7.2 6.6 9.1 5.0
2.5 4.7 2.7 3.5 3.1 3.5
in Eemian d18O, that the differences in elevation between the Camp Century, GRIP, GISP2, NGRIP, and Renland sites during the Eemian were almost equal to those presently observed. Furthermore, the elevated Eemian d18O values at the Dye-3 site indicate that the southern part of the Greenland ice sheet was lowered by some 400 m relative to the rest of the ice sheet (the 1.5% extra change at Dye-3 added to a 1% change due to ice flow induced extra lowering of the Eemian d18O). It is more difficult to access the absolute elevation of the main ice sheet during the Eemian. The general 3% increase in d18O could be attributed both to changes in elevation and changes in temperature, as the Greenland climate during the Eemian was significantly warmer than the present climate. The fact that the small Renland ice cap experienced almost the same increase in d18O as most of the Greenland ice sheet does however put constraints on the possible changes in elevation. The Renland ice cap is only about 330 m thick, and due to its position on a high elevation plateau, it cannot grow any thicker. Furthermore, the fact that a more than 3-m-thick layer of Eemian ice has been found close to bedrock at Renland (see Figure 5) does suggest that the Renland ice cap was well preserved during the Eemian. The ice flow-driven thinning rate in the Renland ice cap estimated by Johnsen and Dansgaard (1992), probably could not have allowed the preservation of a 3-m thick slab of Eemian ice, had the Renland ice cap been significantly thinner during the Eemian than today. It is, therefore, reasonable to assume that the Eemian thickness of the Renland ice cap was not very far from the 330 m observed today, thus imposing narrow constraints on thickness changes at Camp Century, GISP2, GRIP, and NGRIP as well. Given this information, it is obvious that most (if not all) of the 2.5–3.5% Eemian increase in d18O should be attributed to temperature change; not to elevation change. Hence, it is probably only the southern part of the Greenland ice sheet that experienced significant decreases in thickness during the Eemian. In fact it could be argued that a slight increase in the thickness of the northern part of the Greenland ice sheet is supported by the d18O data, as NGRIP and Camp Century Eemian d18O seems to have increased a bit less than the Renland d18O.
See also: Ice Core Methods: Borehole Temperature Records; Chronologies; Overview; Stable Isotopes. Ice Core Records: Antarctic Stable Isotopes; Correlations Between Greenland and Antarctica. Ice Cores: Dynamics of the Greenland Ice Sheet; History of Research, Greenland and Antarctica.
ICE CORE RECORDS | Greenland Stable Isotopes
References Andersen KK, Svensson A, Rasmussen SO, et al. (2006) The Greenland ice core chronology 2005, 15–42 kyr, Part 1: Constructing the time scale. Quaternary Science Reviews 25(23–24): 3246–3257. Shackleton special issue. Basnett T and Parker D (1997) Development of the Global Mean Sea Level Pressure Data Set GMSLP2. UK: Hadley Centre. Climate Research Technical Note 79. Benson CS (1962) Stratigraphic Studies in the Snow and Firn of the Greenland Ice Sheet, p. 63. SIPRE Research Report 70 . Bromwich DH, Toracinta ER, Wei H, Oglesby RJ, Fastook JL, and Hughes TJ (2004) Polar MM5 simulations of the winter climate of the Laurentide Ice Sheet at the LGM. Journal of Climate 17: 3415–3433. Dahl-Jensen D, Mosegaard K, Gundestrup N, et al. (1998) Past temperatures directly from the Greenland ice sheet. Science 282: 268–271. Dansgaard W (1954) The O18-abundance in fresh water. Geochimica et Cosmochimica Acta 6: 241–260. Dansgaard W (1961) The isotopic composition of natural waters. Meddelelser om Grønland 165: 120. Dansgaard W (1964) Stable isotopes in precipitation. Tellus 16: 436–468. Dansgaard W, Clausen HB, Gundestrup N, et al. (1982) A new Greenland deep ice core. Science 218: 1273–1277. Dansgaard W and Johnsen SJ (1969) A flow model and a time scale for the ice core from Camp Century, Greenland. Journal of Glaciology 8: 215–223. Dansgaard W, Johnsen SJ, Clausen HB, et al. (1993) Evidence for general instability of past climate from a 250-kyr ice-core record. Nature 364: 218–220. Dansgaard W and Tauber H (1969) Glacier oxygen-18 content and Pleistocene ocean temperatures. Science 166(3904): 499–502. Epstein S and Sharp RP (1959) Oxygen isotope studies. Transactions, American Geophysical Union 40: 81–84. Herron MM and Langway CC Jr. (1980) Firn densification: An empirical model. Journal of Glaciology 25: 373–385. Johnsen SJ (1977) Stable isotope homogenization of polar firn and ice. In: Proceedings on the Symposium on Isotopes and Impurities in Snow and Ice, pp. 210–219. I.U.G. G. XVI, General Assembly, Grenoble, 1975, Washington, DC. Johnsen SJ, Clausen HB, Cuffey KM, Hoffmann G, Schwander J, and Creyts T (2000) Diffusion of stable isotopes in polar firn and ice: The isotope effect in firn diffusion. In: Hondoh T (ed.) Physics of Ice Core Records, pp. 121–140. Sapporo: Hokkaido University Press. Johnsen SJ, Clausen HB, Dansgaard W, et al. (1992) Irregular glacial interstadials recorded in a new Greenland ice core. Nature 359: 311–313. Johnsen SJ and Dansgaard W (1992) On flow model dating of stable isotope records from Greenland ice cores. NATO ASI Series I 2: 13–24. Johnsen SJ, Dansgaard W, and White JWC (1989) The origin of Arctic precipitation under present and glacial conditions. Tellus 41B: 452–468.
409
Johnson KS, Gordon RM, and Coale KH (1997) What controls dissolved iron in the world ocean?. Marine Chemistry 57: 137–161. Langway CC Jr. (1967) Stratigraphic Analysis of a Deep Ice Core from Greenland, CRREL. Research Report 77, p. 130. Laursen EV (2003) DMI Monthly Climate Data, 1873–2002, Contribution to Nordic Arctic Research Programme (NARP). Technical Report 03–25 from the Danish Meteorological Institute. Meese DA, Gow AJ, Alley RB, et al. (1997) The Greenland Ice Sheet Project 2 depth-age scale: Methods and results. Journal of Geophysical Research 102: 26411–26423. Merlivat L and Jouzel J (1979) Global climatic interpretation of the deuterium-oxygen 18 relationship for precipitation. Journal of Geophysical Research C8: 5029–5033. North Greenland Ice Core Project members (2004) High-resolution record of Northern Hemisphere climate extending into the last interglacial period. Nature 431: 147–151. Rasmussen SO, Andersen KK, Svensson AM, et al. (2006) A new Greenland ice core chronology for the last glacial termination. Journal of Geophysical Research 111: D06102. doi:10.1029/2005JD006079. Rasmussen SO, Vinther BM, Clausen HB, and Andersen KK (2007) Early Holocene oscillations recorded in three Greenland ice cores. Quaternary Science Reviews 26: 1907–1914. Ruddiman WF and McIntyre A (1981) The North Atlantic Ocean during the last deglaciation. Palaeogeography, Palaeoclimatology, Palaeoecology 35: 145–214. Schwander J, Stauffer B, and Sigg A (1988) Air mixing in firn and the age of the air at pore close-off. Annals of Glaciology 10: 141–145. Svensson A, Andersen KK, Bigler M, et al. (2006) The Greenland Ice Core Chronology 2005, 15–42 kyr, Part 2: Comparison to other records. Quaternary Science Reviews 25(23–24): 3258–3267. Shackleton special issue. Vinther BM, Buchardt SL, Clausen HB, et al. (2009) Holocene thinning of the Greenland ice sheet. Nature 461: 385–388. Vinther BM, Clausen HB, Johnsen SJ, et al. (2006) A synchronized dating of three Greenland ice cores throughout the Holocene. Journal of Geophysical Research 111: D13102. doi:10.1029/2005JD006921. Vinther BM, Clausen HB, Johnsen SJ, et al. (2008) Reply to comment on: A synchronized dating of three Greenland ice cores throughout the Holocene. Journal of Geophysical Research 113: D12306. doi:10.1029/2007JD009083. Vinther BM, Johnsen SJ, Andersen KK, Clausen HB, and Hansen AW (2003) NAO signal recorded in the stable isotopes of Greenland ice cores. Geophysical Research Letters 30: 1387. Vinther BM, Jones PD, Briffa KR, et al. (2010) Climatic signals in multiple highly resolved stable isotope records from Greenland. Quaternary Science Reviews 29: 522–538.
The Early Isotope Work By doing stable isotope investigations of frontal precipitation over Copenhagen, Denmark, in the early 1950s, Willi Dansgaard discovered a correlation between condensation temperatures and isotopic ratios. Subsequently, Dansgaard proposed in a groundbreaking paper that paleoclimate could be studied by doing stable isotope measurements on Greenland ice cores (Dansgaard, 1954). At this time, no ice cores were available from Greenland for such investigations, and the immense scientific scope of this proposal was hardly realized even 15 years later when the first results from the Camp Century core became available. Carl Benson of the US Army Snow, Ice, and Permafrost Research Establishment (SIPRE) was, however, already doing pit studies in Northwest and Central Greenland in the early and mid-1950s (Benson, 1962). In cooperation with Epstein and Sharp, he discovered the very strong annual isotopic cycle in the Greenland snow pack. Eventually, this cycle became most useful for measuring accumulation rates and dating ice cores. When the Site 2 ice core (see Figure 1), which was drilled by SIPRE in preparation for the International Geophysical Year’s Antarctic activities in 1957, became available in 1955, Chester C Langway, also from SIPRE, initiated the modern era of multidisciplinary ice-core studies including detailed isotope profiles deep down in the core (Langway, 1967). The Site 2 isotope data (Epstein and Sharp, 1959) showed a clear annual cycle below the firn layers, but smoothed compared with the isotope data from Benson’s pit studies (Langway, 1967). Thus, Langway discovered the effect of firn diffusion discussed below. Willi Dansgaard continued his studies in Greenland and discovered the isotopic latitude effect for Greenland and the strong correlation between mean annual temperatures and isotopic values on the ice cap (Dansgaard, 1961). He also used the Rayleigh condensation model to improve the understanding of the physical basis for the climatic information contained in natural precipitation. Most of his earlier results are summarized in his classic Tellus paper (Dansgaard, 1964) where he also provided the first results from the International Atomic Energy Agency (IAEA) precipitation network and introduced the concept of deuterium excess. The Camp Century deep ice core was recovered by Lyle Hansen and his coworkers from the US Army Cold Region Research and Engineering Laboratories in 1966. Two years later, Dansgaard suggested that Langway do a comprehensive isotope study on the core, a proposal that Langway accepted. This most fruitful cooperation was later extended with the group of Hans Oeschger from Bern, Switzerland, and became the basis for the Greenland Ice Sheet Program (GISP) that managed to recover a deep core from Dye-3 South Greenland with a new Danish deep drill in 1981. When the first Dye-3 isotopic results were published (Dansgaard et al., 1982), the immense impact of the isotopic paleostudies on Greenland
deep ice cores was firmly established. In the early 1970s, the Danish group also built a unique inlet system for running 512 daily mass-spectrometric oxygen-isotope measurements. The original aim was to date the Dye-3 core by counting annual isotopic cycles through most of the Holocene by measuring a total of 80 000 samples. However, this system greatly strengthened the isotope work done with subsequent Greenland ice cores and over one million samples were analyzed using it.
Shallow Ice Cores – The Role of Diffusion A large number of ice cores reaching depths of 100–200 m have been drilled at various locations on the Greenland ice sheet. Such cores offer a time span ranging from a few centuries to thousands of years, depending on core depth and local accumulation rates. As instrumental climatic observations are limited to the past two centuries for most of the Northern Hemisphere, the records of d18O from shallow ice cores are a valuable asset when investigating relationships between climate and d18O. Due to the multiple cores available, it is possible to stack the records and thereby reduce the glaciological noise in the d18O data. Using cores from various locations, it is also possible to resolve local climatic variations affecting only parts of the ice sheet. The investigation of d18O records from the upper part of the ice cores is complicated by water vapor diffusion in the pores of snow and firn in the top 60–70 m of the ice sheet. This diffusion tends to smooth the high-frequency oscillations of the d18O. Figures 2 and 3 show the diffusion-related decay of the annual d18O cycle for a central Greenland shallow core.
The Climate Signal in Seasonal d18O The close relationship between d18O and site temperature makes d18O a valuable proxy both for Greenland climate and Northern Hemisphere atmospheric circulation patterns. To exploit the potential of Greenland d18O records fully Vinther et al. (2003, 2010) have shown that it is essential to consider seasonal oscillations in the d18O records. In the vast areas of the Greenland ice sheet where annual accumulation exceeds 20 cm of ice equivalent, seasonal oscillations of d18O can be retrieved through most of the Holocene. In areas of lower accumulation, the seasonal d18O cycles are obliterated by the diffusion in the upper snow and firn layers. As can be seen in Figures 2 and 3, the seasonal cycle of d18O is influenced by diffusion even in areas with accumulation rates larger than 20-cm ice equivalent per year. In the top 60–70 m of the ice sheet, a marked decay in amplitude of the d18O cycle can be observed. If a climatic interpretation of seasonal d18O values is to be pursued, it is necessary to
403
ICE CORE RECORDS | Greenland Stable Isotopes
understand the diffusion process that masks the seasonal oscillations in the d18O signal.
−90
−80 −70 −60 −50 −40 −30 −20 −10
0
10
80
80
76
76
72
72
68
68
64
64
60
60 −50
−40
−30
d 18O (per mil)
Figure 1 Locations of selected Greenland ice cores. The Site 2 core was the first ice core drilled in Greenland. The Camp Century, Dye-3, GISP2, GRIP, NGRIP, and Renland cores all contain ice from the Holocene, Glacial, and Eemian periods.
In a publication by Johnsen (1977), the first step was taken in comprehending the physics behind the diffusion process. Later papers by Merlivat and Jouzel (1979), Herron and Langway (1980), and Schwander et al. (1988) led to the Johnsen et al. (2000) paper in which a physical model covering both d18O and dD diffusion was presented and validated. The Johnsen et al. (2000) model predictions of the diffusion-driven dampening of the annual d18O cycle can be seen in Figure 3. Using the Johnsen et al. (2000) diffusion model and the mathematical formulation of the diffusion problem presented in Johnsen (1977), it is possible to correct d18O series for diffusion. An example of diffusion-corrected d18O data can be seen in Figure 2(b). Having diffusion-corrected data, the underlying seasonally resolved climatic signal can be retrieved from ice-core d18O records. It is easy to demonstrate the importance of correcting the seasonal d18O records for diffusion before any climatic interpretation is pursued. Four different central Greenland ice cores were drilled in 1984/1985 about 500 km from the coastal Tasiilaq meteorological station. Table 1 demonstrates that correlations between winter season d18O and December-to-March temperatures at the Tasiilaq station (Laursen, 2003) are greatly improved by diffusion correcting the records. Furthermore, it can be seen that the ice-core records from the areas of modest accumulation (Sites E and G) benefit the most from the diffusion correction. Using the approach of diffusion correction on several Greenland ice cores, Vinther et al. (2003, 2010) demonstrated that winter season d18O from Greenland ice cores contained both a distinct Greenland temperature signal and the signature of the North Atlantic Oscillation (NAO), a large-scale atmospheric flow pattern. In fact, Greenland annual average temperatures were found to be better captured by winter season d18O data than annual average d18O data. A linear regression between the multi-ice core winter d18O signal and Northern Hemisphere December-to-March sea level pressure (slp) is seen in Figure 4. The NAO pattern with opposite pressure trends over Iceland and the Azores is clearly seen in Figure 4. The regression was carried out for the years
−28
−28
−30
−30
−32
−32
−34
−34
−36
−36
−38
−38
−40
2
3
4 5 Depth (m)
(a) 18
6
7
70 (b)
70.5
71 71.5 Depth (m)
72
d 18O (per mil)
404
−40 72.5
Figure 2 Measurements of d O from different depths in the Site E central Greenland ice core. Left: near the surface before firn diffusion dampens the annual cycle. Right: below the diffusion zone, the annual cycle is significantly dampened. The thin curve is diffusion-corrected data. Annual accumulation at Site E amounts to 21 cm of ice equivalent.
ICE CORE RECORDS | Greenland Stable Isotopes
405
10 20
Depth (m)
30 40 50 60 70 Observed Modeled
80
0
0.5
1
1.5 2 2.5 Amplitude (per mil)
3
3.5
4
Figure 3 Amplitude of the annual d18O cycle in the Site E central Greenland ice core and model estimate.
Table 1 Correlations (r) of winter d18O with Tasiilaq December-toMarch temperatures (1895–1982) Ice core
Site A
Site B
Site E
Site G
Mean ABEG
Accumulation (ice equivalent; cm) r (before diffusion correction) r (after diffusion correction)
29
31
21
25
–
0.29
0.43
0.25
0.21
0.36
0.33
0.50
0.41
0.37
0.52
1872–1970 and the slp-data are from the HadSLP1 dataset (Basnett and Parker, 1997).
Deep Ice Cores – Climate During 125 Millennia During the past 40 years a total of five deep ice cores have been drilled all the way through the Greenland ice sheet. All five records contain ice from the entire Holocene, the last glaciation, and the Eemian. A sixth ice core drilled through the small east coast Renland ice cap also reaches back into the Eemian (for drill site locations see Figure 1). Measurements of d18O have been carried out on the full length of all six ice-core records. From Figure 5 it is clear that the six d18O records are very similar for the past 40 000 years. Below 40 000 years, the Dye-3 record is disturbed by ice flow, as is the Renland core below 60 000 years. The Greenland Icecore Project (GRIP) and GISP2 records are affected by flow-induced foldings below 105 000 years. The North Greenland Icecore Project (NGRIP) core presents the only d18O record known to be stratigraphically continuous into the Eemian. However, Eemian ice is found in all six ice cores.
Timescales for Deep Ice Cores A prerequisite for the climatic interpretation of the deep ice cores is having a well-established depth–age relation. For most
Figure 4 December–March slp-response to a one-sigma increase in winter d18O in a series based on multiple Greenland ice cores. Contour interval is 0.4 hPa, red contours are positive while blue are negative.
of the Holocene, counting annual layers using the seasonal oscillations in d18O is a very robust way of obtaining a timescale for the Greenland ice cores. During the glacial, generally lower accumulation rates lead to diffusion-caused obliteration of the annual d18O cycles, rendering annual layer counting impossible. The first approach to date the glacial part of the ice cores relied on analytical 2D flow model estimates of the progressive thinning of annual layers (Dansgaard and Johnsen, 1969). The flow model presented by Dansgaard and Johnsen (1969) (D–J model) proved to be very adaptive and Johnsen and Dansgaard (1992) refined the analytical D–J model to account for both increased shear near the bottom of the ice sheet and sliding of the ice sheet over the bedrock. The D–J model can also be modified to include flow due to bottom melting affecting the ice sheet. To include temporal variations of the accumulation rate, a numerical version of the D–J model was required. The first effort to create such a model was made by Johnsen et al. (1992). Numerical D–J flow models have been used to provide timescales to both the GRIP and the NGRIP ice cores (NGRIP members, 2004). These models require past accumulation rates as an input parameter. This requirement can be fulfilled by using the relationships between d18O and accumulation rate at the drill sites. Figure 6 shows the Dye-3, GRIP, and NGRIP d18O profiles on timescales derived from numerical D–J flow models. As the bottom-most part of the Dye-3 core is disturbed by ice flow, only the past 40 000 years of the Dye-3 profile is shown. The similarity of the three profiles is striking. The first effort to count annual layers during the glacial was carried out on the GISP2 core (Meese et al., 1997). Seasonally changing features in the visual stratigraphy of the GISP2 core were used for counting the layers. Surprisingly, the glacial GISP2 timescale derived from this methodology did not show a persistent relationship between the d18O and the
406
ICE CORE RECORDS | Greenland Stable Isotopes
Greenland deep ice core isotopes Camp century
Dye-3
GISP2
GRIP
Renland
NGRIP
305
10
1600 1600
1150
1
1800
1500
1
1
1
1
1800
15
1
1800 2 4
Depth (m)
14
2200
8
4
4
4
8
2200
1900
2400
19
12
14
12 14
315
50 14 2500
19
2600
2600
21
23
23
2800
2800
3000
3000
2000
21
320
3000
−35 −30 −25
−45 −40 −35 −30
−45 −40 −35 −30
60
80
21 23
1350
−45 −40 −35 −30 −25
18 20
21
23
40
12
14
2400
18
1950
1300
8
14
8
30
2000
12
8
12
21
2
20
NGRIP age (ka)
4
12
2
2000
4
8
1250
2
2000
2
1850
1200
2
310
−30 −25
23 24 25
100
120
−45 −40 −35 −30
d18O (‰)
Figure 5 Early Holocene, Glacial, and Eemian d18O records from six Greenland ice cores are shown. The cores are all on depth scales. The NGRIP timescale is shown, and the numbers on the left side of the scale give the most prominent interstadials.
Dye-3, GRIP and NorthGRIP isotopic profiles 1
−30
2
4
GRIP NorthGRIP Dye-3
8
d18O (‰)
Eem
−35
12
14
18
20
21
23
24
25
−40 −45 0
20
40
60
80
100
120 × 103
Age (GRIP2001 years b2k)
Figure 6 Profiles of d18O from the Dye-3, GRIP, and NGRIP cores, on the ss09sea flow-modeled timescale. Numbers are given for the most prominent interstadials. b2k, before AD 2000.
accumulation rate back in time. Such a relation is underpinned both by modern observations and the following simple argument: a change in d18O is associated with a change in site air temperature, and as the air temperature determines the maximum water vapor content of the air, the amount of moisture reaching the site will be related to observed d18O values. Based on a new effort to date the glacial part of the NGRIP core using both seasonal oscillations in chemical impurities and visual stratigraphy, it is now believed that the apparent breakdown of the d18O/accumulation relationship observed in the GISP2 core is an artifact caused by erroneous
interpretation of the GISP2 visible strata (Andersen et al., 2006; Svensson et al., 2006).
The Holocene The major part of the deep ice cores consists of Holocene ice. Looking at Figure 7, it is obvious that millennial scale trends in the ice core records are different during the course of the Holocene. Vinther et al. (2009) have shown that these differences are most likely due to a combination of past Greenland ice sheet elevation change and upstream effects (i.e., that ice in
ICE CORE RECORDS | Greenland Stable Isotopes
407
2
d18O deviation (per mil)
0
-2
Camp century NGRIP GRIP GISP2 DYE-3
-4
Agassiz/Renland
-6
0
2000
4000
6000 Age b2k (years)
8000
10 000
Figure 7 Holocene profiles of d18O deviations with respect to present. Heavy lines show millennial scale trends in the data. All profiles are shown on the GICC05 time scale, except for GISP2 d18O which is shown on the original GISP2 time scale. b2k, before AD 2000 .
the lower strata originates from higher elevation positions upstream from the drill site). Furthermore, Vinther et al. (2009) demonstrated, that the average d18O from the two small ice caps of Renland (eastern Greenland) and Agassiz (Ellesmere Island north west of Greenland) can be used to establish the Holocene climatic d18O evolution, if corrected for post-glacial rebound (the corrected average d18O is shown as the black curve in Figure 7). The differences between the Renland/Agassiz d18O average and d18O from the other five ice cores shown in Figure 7 can thus be converted directly to past changes in depositional elevation at the five locations using the observed d18O-elevation relationship for Greenland. A conversion of the Renland/Agassiz d18O average to past Greenland temperatures show a climatic optimum 2–3 C warmer than present (Vinther et al., 2009) in good agreement with results from Dahl-Jensen et al. (1998) showing that borehole temperature measurements at GRIP and Dye-3 imply that temperatures on the ice sheet were, 3 C higher than present during the climatic optimum. All but one of the d18O records shown in Figure 7 are plotted on the Greenland Ice Core Chronology 2005 (GICC05). GICC05 is a newly counted timescale that synchronizes the ice cores using common volcanic reference horizons (Rasmussen et al., 2006; Vinther et al., 2006, 2008). Exploiting the synchronization of the d18O records (Rasmussen et al., 2007) showed that the only significant common events in the early Holocene Dye-3, GRIP, and NGRIP records are the decreases in d18O some 8200, 9300, and 11 400 years ago. These events are known as the 8.2 ka and 9.3 ka cold events and the Preboreal oscillation, and they are observed in ocean sediment cores, tree ring data, and stalagmite records from all over the Northern Hemisphere. The events are believed to be associated with very large freshwater outbursts from the decaying Laurentide ice sheet. The large amount of freshwaters would temporarily weaken the North Atlantic overturning
circulation thereby diminishing the oceanic heat transport towards the polar region.
The Last Glaciation The abrupt shifts in d18O leading to the warm Greenland interstadials (Dansgaard et al., 1993; Johnsen et al., 1992) are often referred to as Dansgaard–Oeschger events. These events are the most prominent features observed in the glacial parts of all Greenland ice cores (see Figures 5 and 6). The shifts are most probably due to large-scale reorganizations of the atmospheric and oceanic circulations during the glacial. Hence, full glacial (stadial) conditions are probably related to a shutdown of the North Atlantic overturning circulation, while the interstadials are warmed by the return of the overturning circulation. Changes from stadial climatic conditions to the 10–15 C warmer interstadial climate are believed to have happened within a decade or two. The glacial climate is therefore characterized by a degree of instability that cannot be found during the Holocene. This instability is likely to be caused by the N–S shifts of the atmospheric polar front that follows the oceanic polar front (Ruddiman and McIntyre, 1981) in the glacial. The d18O at the polar front is close to 8% Dansgaard and Tauber (1969), which is also likely valid in the glacial. The Greenland d18O values are mainly driven by the latitudinal gradient of 0.54% d18O/lat, the distance to the polar front and elevation changes Dansgaard (1961) and Johnsen et al. (1989). Johnsen et al. (1997) estimated the d18O short-term annual RMS (Root Mean Square) variability from the GRIP ice core (Figure 5 in that paper). The records of annual variability (Figure 5(c)) and the d18O (Figure 5(a)) are in anticorrelation. The RMS variability for Holocene, interstadials, and stadials are 0.9%, 1.8%, and 3.4%, respectively. The increased variability is partly due to lower accumulation rate, in the interstadials and stadials, and partly due to the rapid N–S shifts of the polar front during these cold periods.
408
Table 2
ICE CORE RECORDS | Greenland Stable Isotopes Average values and differences of d18O during different climatic periods (all values are given in %)
Drill site Camp century Dye-3 GISP2 GRIP NGRIP Renland
Present (P) 29.5 27.6 35.0 35.1 35.4 27.3
Stadial (S) 41.9 36.1 42.2 41l.9 44.5 32.1
Johnsen et al. (1992) found that an intercomparison between glacial d18O values observed in different Greenland ice cores shows important differences in the size of the d18O increase associated with the shift from stadial to Holocene climatic conditions. The Camp Century core shows the largest d18O changes while the changes at Renland are the most modest (see Table 2 and Figure 5). It is most likely that the large changes in Camp Century (and NGRIP) d18O are related to elevation changes, but the changes could also be partly due to a strongly depleted northwesterly source of precipitation during the stadials at those sites. Such a source would stem from a split in the polar jet-stream, with branches both north and south of the Laurentide ice sheet. Precipitation arriving from the northwest would not be able to reach the Renland ice cap as it is situated east of the main ice sheet, and therefore, is shielded from any such precipitation. Hence, the relatively modest decrease in Renland d18O values during the stadials can be understood in terms of the Renland ice cap not receiving precipitation from the depleted northern branch of a split jetstream. Recent atmospheric modeling efforts suggesting a split jetstream during the glacial in the winter (i.e., Bromwich et al., 2004) therefore enjoy significant support from observations of stadial d18O in Greenland ice cores. The whole glacial period in Camp Century is 3% colder compared to the present than seen in NGRIP. This difference is most likely caused by 500– 600 m higher elevation at the Camp Century compared to the NGRIP elevation.
The Eemian Ice from the last interglacial, the Eemian, has been found near the bedrock in all deep ice cores retrieved from the Greenland ice sheet (see Figure 5). It is, therefore, safe to conclude that most of Greenland was covered by ice during the Eemian. By using the d18O values observed in Eemian ice, it is possible to further quantify the changes in the size of the Greenland ice sheet during the Eemian. Comparing present day d18O values to those observed in the Eemian ice, it is clear that the Eemian had 2.5–3.5% higher d18O values than presently observed all across Greenland (Table 2 and Figure 5) except for at the southern Dye-3 site. At Dye-3, Eemian d18O was 4.7% higher than recent values. The observed d18O value at a site on the Greenland ice sheet is dependent on elevation (the gradient is approximately 0.6% per 100 m altitude according to Johnsen et al., 1989). It can therefore be inferred from the rather uniform (3%) increase
Eemian (E) 27.0 22.9 32.3 31.6 32.3 23.8
P–S
E–P
12.4 8.5 7.2 6.6 9.1 5.0
2.5 4.7 2.7 3.5 3.1 3.5
in Eemian d18O, that the differences in elevation between the Camp Century, GRIP, GISP2, NGRIP, and Renland sites during the Eemian were almost equal to those presently observed. Furthermore, the elevated Eemian d18O values at the Dye-3 site indicate that the southern part of the Greenland ice sheet was lowered by some 400 m relative to the rest of the ice sheet (the 1.5% extra change at Dye-3 added to a 1% change due to ice flow induced extra lowering of the Eemian d18O). It is more difficult to access the absolute elevation of the main ice sheet during the Eemian. The general 3% increase in d18O could be attributed both to changes in elevation and changes in temperature, as the Greenland climate during the Eemian was significantly warmer than the present climate. The fact that the small Renland ice cap experienced almost the same increase in d18O as most of the Greenland ice sheet does however put constraints on the possible changes in elevation. The Renland ice cap is only about 330 m thick, and due to its position on a high elevation plateau, it cannot grow any thicker. Furthermore, the fact that a more than 3-m-thick layer of Eemian ice has been found close to bedrock at Renland (see Figure 5) does suggest that the Renland ice cap was well preserved during the Eemian. The ice flow-driven thinning rate in the Renland ice cap estimated by Johnsen and Dansgaard (1992), probably could not have allowed the preservation of a 3-m thick slab of Eemian ice, had the Renland ice cap been significantly thinner during the Eemian than today. It is, therefore, reasonable to assume that the Eemian thickness of the Renland ice cap was not very far from the 330 m observed today, thus imposing narrow constraints on thickness changes at Camp Century, GISP2, GRIP, and NGRIP as well. Given this information, it is obvious that most (if not all) of the 2.5–3.5% Eemian increase in d18O should be attributed to temperature change; not to elevation change. Hence, it is probably only the southern part of the Greenland ice sheet that experienced significant decreases in thickness during the Eemian. In fact it could be argued that a slight increase in the thickness of the northern part of the Greenland ice sheet is supported by the d18O data, as NGRIP and Camp Century Eemian d18O seems to have increased a bit less than the Renland d18O.
See also: Ice Core Methods: Borehole Temperature Records; Chronologies; Overview; Stable Isotopes. Ice Core Records: Antarctic Stable Isotopes; Correlations Between Greenland and Antarctica. Ice Cores: Dynamics of the Greenland Ice Sheet; History of Research, Greenland and Antarctica.
ICE CORE RECORDS | Greenland Stable Isotopes
References Andersen KK, Svensson A, Rasmussen SO, et al. (2006) The Greenland ice core chronology 2005, 15–42 kyr, Part 1: Constructing the time scale. Quaternary Science Reviews 25(23–24): 3246–3257. Shackleton special issue. Basnett T and Parker D (1997) Development of the Global Mean Sea Level Pressure Data Set GMSLP2. UK: Hadley Centre. Climate Research Technical Note 79. Benson CS (1962) Stratigraphic Studies in the Snow and Firn of the Greenland Ice Sheet, p. 63. SIPRE Research Report 70 . Bromwich DH, Toracinta ER, Wei H, Oglesby RJ, Fastook JL, and Hughes TJ (2004) Polar MM5 simulations of the winter climate of the Laurentide Ice Sheet at the LGM. Journal of Climate 17: 3415–3433. Dahl-Jensen D, Mosegaard K, Gundestrup N, et al. (1998) Past temperatures directly from the Greenland ice sheet. Science 282: 268–271. Dansgaard W (1954) The O18-abundance in fresh water. Geochimica et Cosmochimica Acta 6: 241–260. Dansgaard W (1961) The isotopic composition of natural waters. Meddelelser om Grønland 165: 120. Dansgaard W (1964) Stable isotopes in precipitation. Tellus 16: 436–468. Dansgaard W, Clausen HB, Gundestrup N, et al. (1982) A new Greenland deep ice core. Science 218: 1273–1277. Dansgaard W and Johnsen SJ (1969) A flow model and a time scale for the ice core from Camp Century, Greenland. Journal of Glaciology 8: 215–223. Dansgaard W, Johnsen SJ, Clausen HB, et al. (1993) Evidence for general instability of past climate from a 250-kyr ice-core record. Nature 364: 218–220. Dansgaard W and Tauber H (1969) Glacier oxygen-18 content and Pleistocene ocean temperatures. Science 166(3904): 499–502. Epstein S and Sharp RP (1959) Oxygen isotope studies. Transactions, American Geophysical Union 40: 81–84. Herron MM and Langway CC Jr. (1980) Firn densification: An empirical model. Journal of Glaciology 25: 373–385. Johnsen SJ (1977) Stable isotope homogenization of polar firn and ice. In: Proceedings on the Symposium on Isotopes and Impurities in Snow and Ice, pp. 210–219. I.U.G. G. XVI, General Assembly, Grenoble, 1975, Washington, DC. Johnsen SJ, Clausen HB, Cuffey KM, Hoffmann G, Schwander J, and Creyts T (2000) Diffusion of stable isotopes in polar firn and ice: The isotope effect in firn diffusion. In: Hondoh T (ed.) Physics of Ice Core Records, pp. 121–140. Sapporo: Hokkaido University Press. Johnsen SJ, Clausen HB, Dansgaard W, et al. (1992) Irregular glacial interstadials recorded in a new Greenland ice core. Nature 359: 311–313. Johnsen SJ and Dansgaard W (1992) On flow model dating of stable isotope records from Greenland ice cores. NATO ASI Series I 2: 13–24. Johnsen SJ, Dansgaard W, and White JWC (1989) The origin of Arctic precipitation under present and glacial conditions. Tellus 41B: 452–468.
409
Johnson KS, Gordon RM, and Coale KH (1997) What controls dissolved iron in the world ocean?. Marine Chemistry 57: 137–161. Langway CC Jr. (1967) Stratigraphic Analysis of a Deep Ice Core from Greenland, CRREL. Research Report 77, p. 130. Laursen EV (2003) DMI Monthly Climate Data, 1873–2002, Contribution to Nordic Arctic Research Programme (NARP). Technical Report 03–25 from the Danish Meteorological Institute. Meese DA, Gow AJ, Alley RB, et al. (1997) The Greenland Ice Sheet Project 2 depth-age scale: Methods and results. Journal of Geophysical Research 102: 26411–26423. Merlivat L and Jouzel J (1979) Global climatic interpretation of the deuterium-oxygen 18 relationship for precipitation. Journal of Geophysical Research C8: 5029–5033. North Greenland Ice Core Project members (2004) High-resolution record of Northern Hemisphere climate extending into the last interglacial period. Nature 431: 147–151. Rasmussen SO, Andersen KK, Svensson AM, et al. (2006) A new Greenland ice core chronology for the last glacial termination. Journal of Geophysical Research 111: D06102. doi:10.1029/2005JD006079. Rasmussen SO, Vinther BM, Clausen HB, and Andersen KK (2007) Early Holocene oscillations recorded in three Greenland ice cores. Quaternary Science Reviews 26: 1907–1914. Ruddiman WF and McIntyre A (1981) The North Atlantic Ocean during the last deglaciation. Palaeogeography, Palaeoclimatology, Palaeoecology 35: 145–214. Schwander J, Stauffer B, and Sigg A (1988) Air mixing in firn and the age of the air at pore close-off. Annals of Glaciology 10: 141–145. Svensson A, Andersen KK, Bigler M, et al. (2006) The Greenland Ice Core Chronology 2005, 15–42 kyr, Part 2: Comparison to other records. Quaternary Science Reviews 25(23–24): 3258–3267. Shackleton special issue. Vinther BM, Buchardt SL, Clausen HB, et al. (2009) Holocene thinning of the Greenland ice sheet. Nature 461: 385–388. Vinther BM, Clausen HB, Johnsen SJ, et al. (2006) A synchronized dating of three Greenland ice cores throughout the Holocene. Journal of Geophysical Research 111: D13102. doi:10.1029/2005JD006921. Vinther BM, Clausen HB, Johnsen SJ, et al. (2008) Reply to comment on: A synchronized dating of three Greenland ice cores throughout the Holocene. Journal of Geophysical Research 113: D12306. doi:10.1029/2007JD009083. Vinther BM, Johnsen SJ, Andersen KK, Clausen HB, and Hansen AW (2003) NAO signal recorded in the stable isotopes of Greenland ice cores. Geophysical Research Letters 30: 1387. Vinther BM, Jones PD, Briffa KR, et al. (2010) Climatic signals in multiple highly resolved stable isotope records from Greenland. Quaternary Science Reviews 29: 522–538.