cp 12 171 2016
Clim. Past, 12, 171–188, 2016
www.clim-past.net/12/171/2016/
doi:10.5194/cp-12-171-2016
© Author(s) 2016. CC Attribution 3.0 License.
Spatial and temporal oxygen isotope variability in northern
Greenland – implications for a new climate record
over the past millennium
S. Weißbach1 , A. Wegner1 , T. Opel2 , H. Oerter1 , B. M. Vinther3 , and S. Kipfstuhl1
1 Alfred-Wegener-Institut,
Helmholtz-Zentrum für Polar- und Meeresforschung, Bremerhaven, Germany
Helmholtz-Zentrum für Polar- und Meeresforschung, Potsdam, Germany
3 Centre for Ice and Climate, Niels Bohr Institute, University of Copenhagen, Copenhagen, Denmark
2 Alfred-Wegener-Institut,
Correspondence to: S. Weißbach ([email protected])
Received: 7 May 2015 – Published in Clim. Past Discuss.: 23 June 2015
Revised: 7 October 2015 – Accepted: 19 January 2016 – Published: 3 February 2016
Abstract. We present for the first time all 12 δ 18 O records
obtained from ice cores drilled in the framework of the North
Greenland Traverse (NGT) between 1993 and 1995 in northern Greenland. The cores cover an area of 680 km × 317 km,
10 % of the Greenland ice sheet. Depending on core length
(100–175 m) and accumulation rate (90–200 kg m−2 a−1 ) the
single records reflect an isotope–temperature history over the
last 500–1100 years.
Lowest δ 18 O mean values occur north of the summit and
east of the main divide as a consequence of Greenland’s topography. In general, ice cores drilled on the main ice divide
show different results than those drilled east of the main ice
divide that might be influenced by secondary regional moisture sources.
A stack of all NGT records and the NGRIP record is
presented with improved signal-to-noise ratio. Compared to
single records, this stack represents the mean δ 18 O signal
for northern Greenland that is interpreted as proxy for temperature. Our northern Greenland δ 18 O stack indicates distinctly enriched δ 18 O values during medieval times, about
AD 1420 ± 20 and from AD 1870 onwards. The period between AD 1420 and AD 1850 has depleted δ 18 O values compared to the average for the entire millennium and represents
the Little Ice Age. The δ 18 O values of the 20th century are
comparable to the medieval period but are lower than that
about AD 1420.
1
Introduction
During the past decades the Arctic has experienced a pronounced warming exceeding that of other regions (e.g.,
Masson-Delmotte et al., 2015). To place this warming in a
historical context, a profound understanding of natural variability in past Arctic climate is essential. To do so, studying climate records is the first step. However, meteorological
measurements in the Arctic are only available for relatively
short time periods; only a few time series start as early as
in the 19th century. Hence, proxy data from climate archives
such as ice cores from the polar ice caps are essential.
Studying the climate of the past centuries allows us to
compare the instrumental data with proxy records and therefore to assess the quality of the proxies for climate reconstructions.
Stable water isotopes (here δ 18 O) in ice cores are commonly used to derive paleotemperatures (e.g., Fischer et al.,
1998c; Johnsen et al., 2000; Steffensen et al., 2008). They
are largely controlled by equilibrium and kinetic fractionation processes during evaporation at the ocean surface, along
the poleward air-mass transport and condensation of precipitation, depending on temperature and moisture conditions
(Dansgaard et al., 1969; Jouzel and Merlivat, 1984; Merlivat
and Jouzel, 1979).
The isotope ratio is not only driven by local temperature
but also affected by several factors like moisture sources
and their proximity to the deposition site, the topography
of the ice sheet and the seasonality of precipitation (Fisher
et al., 1985). In addition the isotope signal is altered by
Published by Copernicus Publications on behalf of the European Geosciences Union.
172
S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
post-depositional processes like wind-induced redistribution
of snow, temperature gradient metamorphism and diffusion
(Johnsen et al., 2000; Pinzer et al., 2012; Steen-Larsen et
al., 2014). Stacked records are used to compensate for effects due to local to regional differences and to improve the
signal-to-noise ratio (Fisher et al., 1985; Masson-Delmotte et
al., 2015; White et al., 1997).
To date, most ice core studies on the Greenland ice sheet
have been carried out point-wise (e.g., Dye 3, GRIP, GISP2,
NGRIP), which begs the question of how representative one
single long ice core record is for deriving a comprehensive
record of past climate. A study of ice cores from southern Greenland revealed that winter season stable water isotopes are largely influenced by the North Atlantic Oscillation (NAO) and are strongly related to southwestern Greenland air temperatures. On the other hand, summer season stable water isotope ratios show higher correlations with North
Atlantic sea surface temperature conditions (Vinther et al.,
2010). In particular, northern Greenland has been little investigated so far. The summit in Greenland’s center is the highest site and separates Greenland into a northern and southern
part. Northern Greenland differs significantly from the south
in terms of lower air temperatures and lower snow accumulation rates (Fischer et al., 1998c). Thus, the results from southern Greenland are not directly transferable to the northern
part.
Northern Greenland’s climate is influenced by different effects than the southern part. One example is the NAO effect,
which is present in the southern and western part of Greenland and is discussed to be reduced in northern Greenland
(Appenzeller et al., 1998).
The cyclones causing the precipitation over northern
Greenland originate in the Baffin Bay and bring dry and
cold air masses from the central Arctic to northern Greenland (Chen et al., 1997).
The dominant westerly winds are blocked by the ice divide, while the northeastern part has very low accumulation
rates below 100 kg m−1 a−1 .
The topographic situation in northern Greenland is special
for δ 18 O studies. In northern Greenland going northward also
means to go downward (lower altitudes).
For a correct estimate of mass balances as well as the response to the ongoing climate change, knowledge of accumulation rates and the spatial distribution of δ 18 O as a temperature proxy is important for the entire Greenland ice sheet.
However, due to northern Greenland’s remoteness its recent
past climate has, up to now, been only scarcely investigated.
Even in the 1990s little was known about northern Greenland. Only few studies had been performed before the Alfred
Wegener Institute’s (AWI) North Greenland Traverse (NGT)
started in 1993. There had been a traverse by Koch and Wegener in 1913 (Koch and Wegener, 1930) and one by Benson in 1952–1953 (Benson, 1962), and there was the British
North Greenland Expedition in 1958 (Bull, 1958), which
studied the accumulation rate in northern Greenland. HowClim. Past, 12, 171–188, 2016
ever, there had been no stable water isotope studies in the
central part of northern Greenland. Fischer et al. (1998c) and
(Schwager, 2000) present the first results from δ 18 O values
of some of the NGT records.
Using the updated accumulation rates of the (compared to
Friedmann et al., 1995; Schwager, 2000) NGT, it was possible to show that the area of lower accumulation rates is much
larger than expected before, which has an influence on the
outlet glaciers (Weißbach et al., 2015).
The NGT ice cores offer, for the first time, the possibility
to study the spatial and temporal variability in stable oxygen isotope records from northern Greenland. Furthermore,
they allow the analysis of the common spatial stable water
isotope signal in northern Greenland by stacking the individual records to significantly reduce the isotopic noise that is
present in a single record due to local peculiarities.
The main objectives of this study are (1) to investigate the
spatial variability in δ 18 O in northern Greenland using this
new set of δ 18 O data and to evaluate the influence of isotopic
noise on a single record, (2) to assess whether stable water
isotope records from sites with low accumulation rates can
be interpreted as climate signals, (3) to present a new robust
stacked δ 18 O record for northern Greenland covering the past
millennium, and (4) to interpret this record in terms of paleoclimate with respect to temporal variability and relation to
large-scale climate information from other proxy records.
2
Material and methods
The ice cores presented here were drilled during the NGT
from 1993 to 1995. In total, 13 ice cores (B16-B23, B26B30) from 12 different sites (Table 1, Fig. 1) were drilled
along the traverse route. The ice cores cover the last 500–
1000 years. The drillings were accompanied by extensive
surface snow studies (e.g., Schwager, 2000).
B21 and B23 as well as B26 to B30 are located on ice
divides (Fig. 1), while B16–B20 were drilled east of the
main ice divide. The NGRIP core (North Greenland Ice Core
Project Members, 2004) was drilled 14.5 km northwest of
B30 following the main ice divide and is therefore included
in this study.
Before analyzing the stable water isotopes, a density profile of each core was measured. To do so, the single core segments (approximately 1 m long) were weighed in the field.
Additional higher-depth-resolution density records were determined using gamma-absorption measurements in the AWI
cold lab (Wilhelms, 1996). Finally, in 2012, density of the
first 70 m of the three cores B19, B22 and B30 was analyzed
by X-ray computer tomography (X-CT; Freitag et al., 2013).
An exponential function fitted to the data taking into account all three types of density data with same respect was
used to calculate water equivalent (w.e.) accumulation rates
and to synchronize the cores.
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S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
173
Table 1. Overview of all NGT drill sites.
Core
ID
B16
B17
B18
B19
B20
B21
B22
B23
B26
B27
B28
B29
B30
Core length
Elevation
Geographic position
(m)
(m a.s.l.)
Latitude
(◦ N)
Longitude
(◦ W)
102.4
100.8
150.2
150.4
150.4
100.6
120.6
150.8
119.7
175.0
70.7
110.5
160.8
3040
2820
2508
2234
2147
2185
2242
2543
2598
2733
2733
2874
2947
73.94
75.25
76.62
78.00
78.83
80.00
79.34
78.00
77.25
76.66
76.66
76.00
75.00
37.63
37.63
36.40
36.40
36.50
41.14
45.91
44.00
49.22
46.82
46.82
43.50
42.00
Selected parts of B30 were also analyzed for electrolytic
conductivity using high-resolution continuous-flow analysis
(Kaufmann et al., 2008).
For the isotopic measurements the ice was cut into samples of 1–5 cm depth resolution, corresponding to 2–10 samples per year. Most of the ice was sampled with 2–2.5 cm
depth resolution. Only at the uppermost parts of the core
were samples cut with lower depth resolution (up to 5 cm).
For some meters of special interest a resolution of 1 cm was
used. After melting, δ 18 O was determined using Delta E and
S mass spectrometers from Finnigan MAT in the AWI laboratory with uncertainties less than 0.1 ‰ as determined from
long-term measurements. Cores B27 and B28 were drilled
at the same site. Parts of core B27 (8.25–11.38 m w.e.), corresponding to AD 1926–1945) were lost, and these were replaced by the record from B28. For the other parts, the mean
of both dated cores was calculated to generate one isotope
record for this site.
Six of the NGT cores (B16, B18, B20, B21, B26 and B29)
were already dated up to a certain depth by annual layer
counting (using density, major ions or δ 18 O) in prior studies (e.g., Fischer and Mieding, 2005; Fischer et al., 1998a,
b; Schwager, 2000). Depending on the availability of data
and differences in snow accumulation rates the dating quality of these cores varies between 1- and 5-year accuracy. For
the other NGT cores annual layer counting was not possible
due to the very low accumulation rates (< 100 kg m−2 a−1 ).
To achieve the same dating quality for all NGT cores for
better comparison and to apply the dating on the whole
core length, we used a new dating procedure for all cores.
From density-corrected (w.e.) high-resolution electrical conductivity profiles (Werner, 1995; Wilhelms, 1996) and SO2−
4 concentration profiles for B16, B18, B21 (Fischer et al.,
1998a, b), B20 (Bigler et al., 2002) and an electrolytic conductivity profile (B30), distinct volcanic horizons were idenwww.clim-past.net/12/171/2016/
Figure 1. Map of Greenland with NGT ice cores (B16–B23, B26–
B30 crosses), deep drilling sites (crosses) and towns (black dots).
The ice surface topography is according to Bamber et al. (2013),
(mapping: Polar Stereographic (WGS84), Standard Parallel 71, Latitude of Projection Origin −39).
tified and used as match points to synchronize the cores (Table 2). Some of the volcanic eruptions show a more pronounced signal in the Greenlandic ice than others. Thus not
all eruptions could be identified in every record.
Between match points, the annual dating was assigned
assuming a constant snow accumulation rate. If a volcanic
match point could not be clearly identified in an ice core, the
next time marker was used to calculate the mean accumulation rate. Below the deepest volcanic match point, the last
calculated accumulation rate was extrapolated until the end
of the core. As the cores were drilled only in the upper part
of the ice sheet (up to 100–175 m depths), layer thinning was
not taken into account.
Clim. Past, 12, 171–188, 2016
174
S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
Table 2. Depth of volcanic horizons used for dating. The given year is the time of aerosol deposition on the Greenland ice sheet. All depths
are given in meter water equivalent. If a horizon could not be clearly identified a dash is shown in the table. A field is empty if the horizon is
deeper than the length of the ice core. The maximum difference is estimated from a comparison between cores dated by annual layer counting
(Mieding, 2005; Schwager, 2000) and the dating used for this study. Also given are the volcanic explosivity index (VEI; Newhall and Self,
1982) and the total Northern Hemisphere stratospheric sulfate aerosol injection (Gao et al., 2008) for each volcanic eruption used.
Year [AD]
Event
1912
Katmai∗
1816
Tambora∗
1783
Laki∗
1739
Tarumai∗
1694
Hekla∗∗
1666
Unknown∗∗
1640
Komagatake∗∗
1601
Huaynaputina∗
1479
Mt. St. Helens∗∗
1259
Samalas∗
1179
Katla∗
934
Eldgjá∗
Max. age of core [AD]
Max. difference [a]
B16
B17
B18
B19
B20
B21
B22
B23
B26
B27
B28
B29
B30
VEI
Sulfate
11.60
24.49
29.36
35.52
–
46.22
49.90
9.31
20.27
24.19
–
34.47
–
–
44.97
58.84
8.62
19.46
23.10
–
32.87
35.93
–
42.95
56.04
11.56
26.17
31.25
–
44.06
48.13
–
–
75.09
9.49
21.54
25.93
–
–
–
–
48.31
–
89.35
98.60
14.27
31.50
37.77
–
–
–
–
69.22
13.69
31.13
37.19
–
–
–
–
68.39
89.42
126.10
14.44
31.91
38.07
11.41
25.97
31.28
–
44.22
48.50
52.36
58.25
76.81
13.12
29.91
35.80
43.07
50.45
–
–
65.94
86.60
122.10
6
7
4
5
4
11.0
58.7
93.0
0
0
0
33.8
46.1
7.4
145.8
0
0
1363
7.38
16.77
19.94
24.10
28.54
31.20
33.69
–
–
68.03
–
99.20
753
7.86
17.27
20.32
24.62
29.16
32.10
34.80
38.70
51.31
72.86
80.04
1470
7
8.48
18.91
22.45
26.90
31.84
34.75
37.48
41.62
54.42
76.60
–
109.20
874
3
775
8
1372
6
1372
1023
1505
4
1195
1763
1471
3
1242
4
4
∗ Sigl et al. (2013); ∗∗ Friedmann et al. (1995).
3
3.1
Results and discussion
Depth–age models and snow accumulation rates
The last millennium was a volcanically active time (Sigl et
al., 2013). The volcanic aerosols deposited on the Greenland
ice sheet can be used as time markers. The depths of peaks
in conductivity and sulfate concentration attributed to certain
volcanic horizons are given in Table 2 as used for our dating
approach.
During the last 500 years, the time period between two detectable eruptions at NGT sites does not exceed 100 years
for any of the cores. This leads to a dating uncertainty for
each core of smaller than 10 years compared to the annually
counted timescales (Mieding, 2005; Schwager, 2000), which
is minimal at the matching points. The three youngest volcanic reference horizons (Katmai, Tambora and Laki) and the
eruptions from AD 1257 (Samalas; Lavigne et al., 2013) and
AD 934 (Eldgjá) were found in all cores, whereas the other
eruptions could not be clearly identified in every ice core.
We could not find a common pattern (e.g., distance, strength
of the eruption) regarding whether or not volcanic horizons
could be observed in all records.
This already indicates a high spatial variability within the
study region related to significant influences of local to regional peculiarities (e.g., wind drift or sastrugi formation).
An overview of the resulting mean accumulation rates for
the entire core lengths for all NGT drilling sites, as well
as the respective ranges, is given in Table 3. According to
our dating, the cores reaching furthest back in time are B18,
B19 and B20, covering more than the last 1000 years. These
northeasterly cores have the lowest accumulation rates with
values below 100 kg m−2 a−1 (B19: 94 kg m−2 a−1 ; B20:
98 kg m−2 a−1 ), whereas the highest mean accumulation rate
is found for B27/28 in the southwest of our study region with
Clim. Past, 12, 171–188, 2016
180 kg m−2 a−1 . Generally, the accumulation rate decreases
from the sites located on the main ice divide in the southwest
of the study area to the northeast.
The observed range of accumulation at one single site is
highest for the southwestern cores (B30 and B29) ranging
between 137 and 161 kg m−2 a−1 (B29). Lowest values are
found for the cores east of the main ice divide (e.g., B17, B18
and B19) ranging between 113 and 119 kg m−2 a−1 (B17).
The length of the records varies depending on accumulation rate and total length of the core. The longest records
are from B19 (back to AD 753) and B20 (back to AD 775).
The following comparisons of the individual records refer to
the longest common time frame (AD 1505–1953). Although
diffusion is known to change isotopic values in the snow, in
this study the data were not corrected for diffusion effects.
While diffusion length is in the range of annual layer thickness, diffusion might be affecting the absolute difference in
isotope content of neighboring years, but the mean over 11
or 30 years will not be affected.
3.2
Regional variability in δ 18 O in northern Greenland
Annual mean δ 18 O records of the NGT cores are displayed
in Fig. 2. Table 4 summarizes the main δ 18 O characteristics
of each core.
The lowest mean δ 18 O values (∼ −37 ‰) in northern
Greenland (B16–B18) and possibly the lowest in Greenland
are found east of the main ice divide and north of the summit, but not at the summit as might be expected. Also, the
lowest firn temperatures were measured at B160–B18 (Table 4). This is in contrast to the findings of Ohmura (1987),
who suggested temperatures similar to the summit for this
region.
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S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
175
Figure 2. Annual δ 18 O records at the 12 NGT sites (this study) and NGRIP (Vinther et al., 2006b), GRIP (Vinther et al., 2010) and GISP2
(Grootes and Stuiver, 1997). Blue values are those below the mean over their common time frame (AD 1505–1953) and red are the higher
ones. Dark-green vertical lines mark the volcanic eruptions (years given at top) used as time markers.
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Clim. Past, 12, 171–188, 2016
176
S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
Table 3. Resulting mean accumulation rates (from the surface to the deepest volcanic horizon and in brackets for their common time frame
(AD 1505–1953)) for each NGT drill site, the lowest and highest rate within the whole core length, the time period from surface to the
deepest volcanic horizon, and the age at the bottom of the ice core calculated by extrapolation of the deepest calculated accumulation rate.
Core
Mean accumulation rate
[kg m−2 a−1 ]
Accumulation rate range
[kg m−2 a−1 ]
Time period
[AD]
Age at bottom of core
[AD]
141 [141]
114 [114]
103 [106]
94 [94]
98 [100]
109 [109]
145 [145]
121 [122]
176 [176]
180 [181]
149 [150]
166 [169]
134–148
113–119
100–110
90–99
90–105
105–113
141–154
116–132
172–190
165–187
137–161
158–178
1640–1993
1479–1993
934–1993
934–1993
1179–1994
1479–1994
1479–1994
1179–1994
1601–1995
1783–1995
1479–1995
1259–1995
1470
1363
874
753
775
1372
1372
1023
1505
1195
1471
1242
B16
B17
B18
B19
B20
B21
B22
B23
B26
B27/B28
B29
B30
Table 4. The 15 m firn temperature, mean annual δ 18 O values for each ice core, the range of the highest and lowest δ 18 O values and the year
they occurred as well as the standard deviation (SD) all given for their common time frame (AD 1953–1505).
Core
B16
B17
B18
B19
B20
B21
B22
B23
B26
B27/B28
B29
B30
NGRIP∗
15 m firn
temperature [◦ C]
Mean δ 18 O
[‰]
δ 18 O range
[‰]
Years [AD] of
lowest/highest value
SD δ 18 O
[‰]
Time period [AD]
(whole core length)
−32.5∗∗∗
−32.3∗∗∗
−32.3∗∗∗
−30.9 (±0.5)∗∗
−30.4∗∗∗
−30.1∗∗∗
−29.8∗∗∗
−29.3 (±0.5)∗∗
−30.3∗∗∗
−30.6 (±0.5)∗∗
−31.6 (±0.5)∗∗
−31.8 (±0.5)∗∗
−37.07
−37.13
−36.53
−35.49
−35.41
−34.53
−34.54
−35.98
−33.86
−34.47
−35.65
−35.46
−35.42
−40.64 to −33.11
−40.06 to −33.89
−41.52 to −31.75
−38.97 to −31.77
−39.34 to −30.69
−38.29 to −30.95
−39.11 to −29.84
−42.11 to −32.23
−37.22 to −29.42
−38.26 to −30.58
−39.22 to −31.59
−38.53 to −31.52
−40.12 to −30.81
1839/1937
1835/1879
1761/1889
1575/1826
1699/1929
1814/1871
1921/1953
1918/1928
1597/1893
1566/1892
1834/1928
1862/1928
1836/1928
0.99
1.08
1.44
1.32
1.42
1.29
1.34
1.28
1.25
1.25
1.18
1.09
1.24
1470–1993
1363-1993
874–1993
753–1953
775–1994
1372-1994
1372–1994
1023–1994
1505–1995
1195–1995
1471–1995
1242–1988
∗ Vinther et al. (2006b); ∗∗ Schwager (2000); ∗∗∗ interpolated from Schwager (2000).
Generally, the cores located east of the main ice divide
show lower mean δ 18 O values than those located on the ice
divide (Fig. 3a). For instance, B29 and B30 are at similar
altitudes and latitudes to B16 and B17 but show significantly
higher values (Fig. 3a).
Figure 4 indicates that accumulation, latitude and altitude
may have a minor impact on the δ 18 O values here. One possible explanation would be additional moisture isotopically
depleted during the transport from directions farther north.
The cores more to the north (B19–B22) were drilled at
lower altitude and therefore record different climate signals
(i.e., from lower air masses) compared to the high-altitude ice
cores that, in turn, record a more smoothed signal of higher
atmospheric layers. Similar effects were observed, for exam-
Clim. Past, 12, 171–188, 2016
ple, in Svalbard (Isaksson et al., 2005), even though at considerably lower altitudes compared to Greenland.
The maximum difference in mean δ 18 O values of individual ice cores is 3.3 ‰ (highest mean δ 18 O in B26: −33.86 ‰;
lowest mean δ 18 O in B17: −37.13 ‰). The standard deviation (SD) for the annual mean values within each core in
the common time frame (AD 1505–1953) is lowest for B16
(0.99 ‰) and highest for B18 (1.44 ‰). We found no general
relation between accumulation rate and standard deviation
of the δ 18 O values for all individual cores, even though the
northern cores with generally lower accumulation rates show
higher standard deviations than the southern cores.
The correlation coefficients between the annual δ 18 O
records of individual ice cores are relatively small (r = 0.1
to 0.36, p < 0.05). This can be partly explained by the fact
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S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
177
Figure 3. Spatial distribution of δ 18 O values in northern Greenland. (a) The mean δ 18 O values of the northern Greenland ice cores in their
common time frame (AD 1505–1953) are shown with color-coded squares. Blue colors represent lower values and red colors higher values.
Mapped mean anomalies of δ 18 O compared to (a) for two different periods: (b) AD 1410–1430 and (c) AD 1920–1940. If a record does not
cover the required time period the square is filled in black.
that the 13 northern Greenland (NG) drilling sites (12 NGT
and NGRIP) are up to 680 km apart from each other. In other
studies where correlated cores are drilled closer together, at
one drill site, higher correlation coefficients were found (e.g.,
at GRIP r = 0.41–0.55 – White et al., 1997; or at NEEM
r ∼ 0.54 – Masson-Delmotte et al., 2015). The strongest correlations are found for the cores from the southwest (B26–
B30) and the lowest for those from the northeast (B19, B20).
There is a significant linear relationship between the distance
between the core sites and their annual δ 18 O correlation coefficient (r = −0.44, p < 0.05). However, it is not always true
that the cores with smallest distance between them have the
highest correlations.
For smoothed values (11-year running mean) the correlation coefficients between the δ 18 O records are only slightly
higher. Only 50 % of the combinations have coefficients
higher than 0.3, and 14 % are lower than 0.1. This indicates
an important influence of regional site-to-site differences.
Variability in δ 18 O is dependent on local (e.g., wind), regional (e.g., position on the ice sheet) and large-scale (e.g.,
circulation patterns) processes. Even adjacent cores may differ considerably according to snow drift (Fisher et al., 1985).
One further reason for the rather low correlations may be attributed to dating uncertainties.
From Fig. 2, we compare our individual NGT δ 18 O records
to other published central to northern Greenland (GRIP,
GISP2, NGRIP) δ 18 O time series. Prominent decadal-scale
maxima and minima occurred mostly isochronally. However,
specific events such as warm periods around AD 1420 or
AD 1920–1930 or a cold period in the 17th century are more
pronounced in the NGT cores compared to summit records.
In Fig. 2 is also obvious that some records show faster
changes between warmer and colder events (e.g., GRIP, B30
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and B26), while others (e.g., B17–B21) remain longer at values higher or lower than their mean (Fig. 2). The longest
warm period (compared to the mean of whole core length)
is found in B19 (with 37 subsequent years warmer than the
mean), while B17 has the longest cold period (28 subsequent
years colder than mean). GRIP, B26, and B27/28 show a
higher frequency with a maximum of about 10 subsequent
warmer or colder years. A frequency analysis of 11-year running mean smoothed data supports these findings. B18–B21
and B29 show much longer main periods (117–248 a) than
B16–B17 and B22–B30 (besides B29, 81–39 a).
In general, the first half of the last millennium was characterized by longer warm or cold anomalies than the second
half and records with more rapid fluctuations are from the
summit and the main ice divide, while those cores drilled
east of the divide have longer periods of positive or negative
anomalies. We conclude that, east of the divide, the climate
conditions are not as variable and therefore the annual δ 18 O
signal is of greater persistence.
The east-to-west difference is also expressed by the dependency of δ 18 O values on longitude (Fig. 4). This is in
line with results from Box (2002), who found that there is
often an opposite trend in air temperatures in east and west
Greenland. The antiphase of temperature records from east
and west Greenland is possibly explained by the importance
of different weather regimes (e.g., Ortega et al., 2014).
The range in δ 18 O in the different cores is different, too.
Cores drilled in the northeast that are characterized by the
lowest accumulation rates have the highest standard deviations (SD) in δ 18 O, which can be partly explained by the fact
that a smaller number of accumulation events scatter more
easily.
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178
S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
Figure 4. Mean δ 18 O (AD 1505–1953) as a function of (a) alti-
tude, (b) latitude, (c) longitude and (d) accumulation rate of northern Greenland ice core drill sites. Cores with higher accumulation
rates (> 145 kg m−2 a−1 ) are given as black dots and lower rates
as black triangles, which is similar to the differentiation between
east of and on the main ice divide. For statistically significant correlations the lines give the linear regression functions (black: mean;
green: higher accumulation rates; blue: lower accumulation rates).
We investigated the relationship between the altitude, latitude and longitude of the drilling sites and the mean δ 18 O values (Fig. 4a, b, c), which are, when considering all records,
statistically significant (p < 0.05) only for longitude and latitude. Regarding their snow accumulation rate we differentiate between two groups: (i) cores with accumulation rates
lower than 145 kg m−2 a−1 mainly located east of the main
ice divide (B16–B21 and B23) and (ii) cores with higher
accumulation rates (B22, B26–B30 and NGRIP). We find
higher δ 18 O ratios for sites with higher accumulation rates
(Fig. 4d). The relationship is weak but becomes stronger for
higher accumulation rates. Buchardt et al. (2012) noted that
the relationship between accumulation rate and δ 18 O is not
distinct for Greenland. Furthermore, Buchardt et al. (2012)
found that the sensitivity of δ 18 O changes to accumulation
rate is smallest in northeastern Greenland (North Central and
North 1972), which is in agreement with our findings.
Among the factors influencing the mean isotopic composition, longitude has the strongest impact (R 2 = 0.56), which
becomes clearest when looking only at the data of group I
(R 2 = 0.93). Figure 4c shows the clear east-to-west gradient
in the mean δ 18 O values in northern Greenland.
If separating between group I (“East”) and group II (“Divide”) there is a strong altitude effect (R 2 = 0.93 and 0.78)
in the data, too.
Clim. Past, 12, 171–188, 2016
These patterns may be explained by different atmospheric
circulation conditions allowing additional moisture from
other sources to reach the region east of the ice divide. This
is supported by the finding of Friedmann et al. (1995), who
suppose, based on data from B16 to B19, that northeastern
Greenland receives more moisture from local sources as the
Greenland Sea, Atlantic Ocean and the Canadian wetlands,
in particular during summer.
We found lower δ 18 O values in the southern and eastern
part of northern Greenland in contrast to the general ideas
of Dansgaard (1954), who expected lower values northward.
That we do not find the lowest values north is a consequence
of different factors in northern Greenland that balance each
other out. More to the north, where we would expect lower
δ 18 O values, the altitude in northern Greenland is decreasing,
which causes higher δ 18 O values (Fig. 1). A multiple linear
regression becomes necessary, as Johnsen et al. (1989) did
before.
Applying this approach to our data, we find
δ(δ 18 O)/(δ(latitude) = −0.30 (±0.40) ‰ degree−1 and
δ(δ 18 O)/δ(altitude) = −0.0035 (±0.0024) ‰ m−1 . The
regression residuals are linearly related to longitude as well
as accumulation rate.
In general, we found correlations with altitude, latitude
and longitude, but the balancing-out effects because of the
special topography in northern Greenland have to be taken
into account.
To study the regional-scale patterns of common variability
in all annual δ 18 O records, we performed a principal component analysis (PCA). All calculations are done for the largest
common time frame of all cores (AD 1505–1953). Other time
periods were used as well, and they show similar results.
Only the first two principal components (PC1 and PC2) are
above the noise level. The first two eigenvectors of the isotopic time series explain 34.1 % of the total variance (PC1:
21.8 %; PC2: 12.3 %). PC1 is similar to the mean of all
records (r = 0.97, p 0.01). It was not possible to assign
PC2 to any climatic relevant signal. The other PCs are dominant in one or two records but are not significant for the total
variance of the entire data set. The loading patterns show a
homogeneous pattern for EOF1 and a bipolar (west–east) result for EOF2.
To summarize, the spatial differences in mean δ 18 O values in northern Greenland can be largely explained by the
influence of the topography of the ice sheet on the regional
climate system. The main ice divide influences the pathways
of air masses, causing lower accumulation rates in the east.
We assume that the temporal variability in a stacked NG
δ 18 O record represents past temperature development.
3.3
The northern Greenland δ 18 O stack and its
paleoclimatic significance
Stable water isotope ratios in ice are widely used as a
proxy for air temperature (Dansgaard, 1964; Johnsen et al.,
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S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
1995; Jouzel et al., 1997b). The comparison to direct airtemperature observation data and proxy data allows for assessment of the quality of the proxy in terms of paleoclimatological interpretation.
To reduce the noise in the single δ 18 O records, we calculated a stacked record by averaging the 13 annual NG δ 18 O
records in their overlapping time periods (NG stack, Fig. 5).
Before stacking, all records were centered and normalized
regarding their common time frame (AD 1505–1953). The
SD of the NG stack (0.44 for AD 1505–1953) is less than
half of the SD in annual δ 18 O records of the individual cores.
Vinther et al. (2010) also point out that stacking is important
to improve the signal-to-noise ratio in areas with low accumulation rate. Local drift noise accounts for half of the total
variance in single-site annual series (Fisher et al., 1985). As
the NG stack before AD 1000 is based on only four records
(< 25 % of the total core numbers), we decided to focus in the
following only on the time period after AD 1000.
As the NG stack is a result of 13 ice cores over a large
area, we assume it is regionally representative.
To investigate the relationship of the NG stack with air
temperature, we used monthly meteorological observations
from coastal southwestern Greenland sites and Stykkishólmur in northwestern Iceland available from the Danish Meteorological Institute (DMI; http://www.dmi.dk; AD 1784–
1993) and the Icelandic Met Office (http://en.vedur.is/; back
to AD 1830), respectively. We selected only the Greenlandic
temperature records longer than 200 years for our study even
though they are at a large distance to the NGT drill sites
(706–2206 km).
The correlation coefficients between the NG stack and
these air-temperature records are shown in Table 5. Dating uncertainties are taken into account by comparing 5year running means. The NG stack shows low but significant
(p < 0.001) correlations with the air temperatures at all sites
(Table 5).
The strongest correlation with annual mean temperature
was found for the merged station data at Greenland’s southeast coast (r = 0.51), and the temperature reconstruction
for the North Atlantic Arctic boundary region of Wood et
al. (2010) (r = 0.55); the lowest was also found for Qaqortoq
(r = 0.39) in the south of Greenland (Table 5). For Stykkishólmur the correlation is in the range of the Greenlandic
ones (r = 0.41). Slightly higher correlations are obtained by
comparing the NG stack to seasonal data. Except for Ilulissat, winter months (DJF) show weaker correlations; spring
(MAM) and summer (JJA) months show stronger correlations with the NG stack.
Comparably low correlations between annual δ 18 O means
and measured temperatures from coastal stations are also reported for the NEEM record (Steen-Larsen et al., 2011).
However, the rather low correlation coefficients might underestimate the real regional δ 18 O–temperature relations because of different reasons.
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179
We expect that the most important reasons are the large
distances and the difference in altitude (i.e., more than
2000 m) between drill sites and the meteorological stations,
which cause them to receive different atmospheric signals.
The stations are located at the coast and are in turn also likely
influenced by local factors such as the occurrence of sea ice.
One other aspect might be seasonality, as argued by SteenLarsen et al. (2011) for the NEEM site. The snowfall in
northern Greenland may be unevenly distributed seasonally.
However, it is not possible to generate sub-annual data for
northern Greenland ice cores due to low accumulation rates.
We find a tendency towards stronger correlation between the
annual δ 18 O and summer (JJA, r = 0.35–0.51) and spring
(MAM, r = 0.36–0.62) temperatures for most of the stations.
This points to a higher proportion of summer snow in the
annual accumulation in northern Greenland, too. SON has
slightly weaker correlation coefficients (r = 0.31–0.5), while
DJF is only significant for Ilulissat and the merged southern
station.
In addition, regional noise factors such as wind drift and
sastrugi formation as well as uncertainties in ice core dating
and the usage of very old observation data have to be taken
into account.
In summary, we consider the northern Greenland δ 18 O
stacked record as a reliable proxy for annual temperature
for northern Greenland. The regional representativeness of
the NG stack is supported by the general similarity to the
NEEM δ 18 O record (Masson-Delmotte et al., 2015) for the
period AD 1724–1994. We found a strong correlation between both records (r = 0.83 for 30-year running mean).
Even single events such as the highest values in AD 1928
and the AD 1810–1830 cooling occur in both records.
Although the NG stack record shows some correlation
with temperature data from coastal Greenland sites, it remains an open question as to how the NG stack δ 18 O variations can be converted into absolute temperature changes
within northeastern Greenland during the last millennium. In
the past, such conversion of isotopic time series of Greenland ice cores was based on a modern analogue approach
taking the observed spatial isotope–temperature gradient of
0.67 ± 0.2 ‰ ◦ C−1 (Dansgaard, 1964; Johnsen et al., 1989)
as a valid calibration for converting isotope records of Greenland ice cores into temperature changes (e.g., Grootes et
al., 1993). The strong confidence of glaciologists in this approach came principally from two observations. (1) Over
both polar ice sheets, the spatial correlation between modern isotope and annual mean temperature is very high and
significant. (2) This empirical observation was theoretically
understood as a consequence of a Rayleigh rainout system
controlling the isotopic composition of meteoric water.
However, for the Greenland area this long-accepted approach has been challenged during the last decade. Two entirely independent analytic techniques, one based on the numerical interpretation of borehole temperatures (e.g., DahlJensen et al., 1998) and the other based on the occlusion proClim. Past, 12, 171–188, 2016
180
S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
Figure 5. Top two panels: the number of cores used for the stack and the standard deviation (SD; gray: annual values; black: 30-year running
mean) of all times. Middle: annual stacked δ 18 O (gray) and smoothed record (30-year running mean). Values more enriched compared to
the mean (AD 1953–1505) are shown in red, while values less enriched are shown in blue. Known climate anomalies are marked: Medieval
Climate Anomaly (MCA, AD 950–1250; Mann et al., 2009), the Little Ice Age (LIA, AD 1400–1900; Mann et al., 1998), and early twentieth
century warming (ETCW, 1920–1940; Semenov and Latif, 2012; Wood and Overland, 2010). Bottom two panels: 30-year running mean on
z levels (centered and normalized data) of stacked northern Greenland δ 18 O records over the last 1000 years. Stack “East”: B16, B17, B18,
B19, B20, B21 and B23. Stack “Divide”: B22, B26, B27, B29, B30 and NGRIP. Values in red are more enriched compared to the mean over
their last 1000 years, and those in blue are less enriched. Also given is the correlation coefficient of 30-year running means between the NG
stack and the sub-stacks (AD 1505–1993). The coefficient for a similar correlation between the two sub-stacks is calculated with r = 0.71.
cess of gases into the ice (e.g., Buizert et al., 2014; Severinghaus et al., 1998), allow for a direct temperature reconstruction at least for some periods of the past. In agreement with
this, both methods point to much lower temporal δ 18 O T−1
slopes ranging between 0.4 and 0.3 ‰ ◦ C−1 (Jouzel et al.,
1997a). Consequently, they indicate a much higher temperature variability in Greenland during the last glacial period.
For the period of the last 9000 years the Greenland average Holocene isotope–temperature relationship has been estimated to be 0.44–0.53 ‰ ◦ C−1 , again substantially lower
compared to the modern spatial gradient (Vinther et al.,
2009). However, as all these studies cover much longer time
periods as compared to our NG stack records, no firm conclusion can be drawn from these studies about an appropriate
isotope–temperature relationship for the last millennium.
Along the NGT firn, temperature measurements at about
15 m depth have been done (Table 4). However, due to their
small range of about 2 K difference, it is difficult to reassess
the general Greenland isotope–temperature relationship from
Johnsen et al. (1989) from the NGT data alone. Schwager (2000) added data from Dansgaard et al. (1969) from
along the EGIG traverse, which was also used in Johnsen et
Clim. Past, 12, 171–188, 2016
al. (1989), to expand the temperature range to derive a more
reliable isotope–temperature gradient. This calculated gradient of 0.7 ± 0.2 ‰ ◦ C−1 is within the gradient uncertainty
range given by Johnsen et al. (1989). Using our updated NGT
data set we get the same results.
If we apply the spatial isotope–temperature gradient of
0.7 ‰ ◦ C−1 from Schwager (2000) for the range of isotope
variations (−1.4 to 2.5 ‰) of the NG stack record, the isotope data translate into temperature changes of −2.0 to 3.6◦ C
(5.6 K) within the last millennium. However, applying instead a temporal gradient of 0.48 ‰ ◦ C−1 as suggested by
Vinther et al. (2009) results in possible temperature changes
of −2.9 to 5.2 ◦ C (8.1 K) within the last 1000 years. Using the most recent temporal glacial–interglacial isotope–
temperature gradients reported by Buizert et al. (2014)
would result in comparable temperature changes. If using
the NEEM gradient of 1.1 ± 0.2 ‰ ◦ C−1 (Masson-Delmotte
et al., 2015), which is valid for AD 2007–1979 in the area
of NEEM, the resulting temperature range of the NG stack
is, at −1.3 to 2.3 ◦ C (3.6 K), a bit smaller than compared to
the Johnsen or Schwager gradient. Nevertheless, the resulting temperature ranges are larger than expected (e.g., Dahl-
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S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
181
Table 5. Correlation coefficients (r) of the stacked δ 18 O record with annual and seasonal (DJF, MAM, JJA and SON) extended Greenland
temperature records ∗ (Vinther et al., 2006a), northwestern Iceland instrumental data∗∗ (Hanna et al., 2004; Jónsson, 1989), annual mean
Greenland ice sheet near-surface air temperatures from combined instrumental and model output ∗∗∗ (Box et al., 2009) and Arctic temperature
reconstruction ∗∗∗∗ (Wood et al., 2010). All correlations are done with 5-year running means and are significant at the 95 % level (p < 0.05).
rannual
Merged South (Greenland)∗
0.51
Ilulissat (Greenland)∗
0.46
Nuuk (Greenland)∗
0.41
Qaqortoq (Greenland)∗
0.39
Stykkishólmur (Iceland)∗∗
0.41
Mean Greenland surface air temperature∗∗∗
Extended instrumental temperature record∗∗∗∗
0.50
0.55
Jensen et al., 1998), which is an additional argument for not
calculating absolute temperatures from the NG stack with the
given gradients.
We conclude that any conversion of the NG stack isotope record into absolute temperature variations during the
last millennium is highly uncertain. Thus, for the following part of the manuscript, we will refer to NG stack isotope anomalies as relative temperature changes in terms of
“warmer” (i.e., isotopically enriched) and “colder” (isotopically depleted) only, but will refrain from converting our ice
core data into absolute temperature changes.
To assess regional differences within northern Greenland,
stacks of subsets of cores will be discussed in terms of interpretation as a temperature proxy. As illustrated in Fig. 4,
we differentiate between two different types of cores: cores
drilled on the ice divide and cores drilled east of the ice divide. Accordingly, in Fig. 5 the overall northern Greenland
δ 18 O stack used in this study is compared to a stack of the
cores of lower accumulation rate drilled east of the main
ice divide (B16, B17, B18, B19, B20, B21 and B23) (stack
“East”) and a stack of those on the ice divide (B22, B26, B27,
B29, B30 and NGRIP) (stack “Divide”) (Fig. 5).
Even though there is a similar overall trend, the three
records show differences in amplitude and timing of warm
and cool events. The correlation between the two sub-stacks
is rather low (r = 0.71 of 30-year running means). In the 11th
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rseasonal
DJF
MAM
JJA
SON
DJF
MAM
JJA
SON
DJF
MAM
JJA
SON
DJF
MAM
JJA
SON
DJF
MAM
JJA
SON
0.47
0.62
0.51
0.5
0.50
0.42
0.36
0.45
–
0.55
0.47
0.47
–
0.50
0.50
0.36
–
0.36
0.35
0.31
Years of overlap
1784–1994
1784–1994
1784–1994
1784–1994
1830–1994
1840–1994
1802–1994
and 12th centuries, we observe a quasi-anti-correlation between stack “East” and stack “Divide”. Even during wellknown climate events such as the Medieval Climate Anomaly
(MCA, AD 950–1250; Mann et al., 2009), the Little Ice Age
(LIA, 1400–1900; Mann et al., 1998) and the early twentieth century warming (ETCW, AD 1920–1940; Semenov and
Latif, 2012), there are significantly different δ 18 O patterns.
For example, stack “Divide” shows colder temperatures during AD 1000–1200. Also, during the 16th century we notice
substantial differences between the two sub-stacks. In stack
“East” events like the AD 1420 or the first part of the LIA
show a higher amplitude.
Stack “East” has a higher correlation to the total NG
stack (r = 0.96) compared to stack “Divide” (r = 0.68) for
the period AD 994–1994. When looking at the time period
AD 1505–1993 with a high number of cores included in both
sub-stacks, it can be seen that the correlation coefficients to
the total NG stack are almost equal (stack “East”: r = 0.95;
stack “Divide”: r = 0.90, p < 0.1). Here, both records reflect
the mean changes in δ 18 O for northern Greenland. Differences before AD 1505 may be artifacts of low core numbers
even though regional differences in climate conditions cannot be ruled out.
We consider the NG stack to be a climate record that displays the overall climate variation independent of local influences such as topography or accumulation rate. In conClim. Past, 12, 171–188, 2016
182
S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
trast, results from studies with only one record become less
spatially representative, as they may be affected by a lower
signal-to-noise ratio and a higher influence of other local
non-climate effects.
3.4
Last millennium climate from a stacked NG δ 18 O
record in relation to other proxy records and
possible forcing factors
The NG stack covers the time between AD 753 and AD 1994
(Fig. 5). For a better visualization of decadal- to centennialscale variability, a 30-year running mean is added. The running mean shows the warmest period around AD 1420 and
the coldest at about AD 1680. The isotopically warmest single year during the last 1000 years in northern Greenland was
AD 1928, whereas AD 1835 was the coldest.
Distinct decadal- to centennial-scale warm and cold
anomalies can be detected in the stacked (Fig. 5) as well
as individual δ 18 O records (Fig. 2) and coincide with wellknown climate anomalies (MCA, LIA, ETCW, marked in
Fig. 5).
We find a pronounced warm period from AD 850 to 1100
which has its maximum between AD 900 and 1000. This is
about 100 years earlier than the described MCA in Mann et
al. (2009). The warm period is followed by a quasi-periodical
change of warm and cold phases observed approximately every 60 to 80 years until about AD 1600. During this phase,
the most distinct warm period is observed around AD 1420.
A longer period of cold temperatures occurred during the
17th and early 19th century and has already been attributed to
the LIA by a prior NGT study that used only four cores (B16,
B18, B21 and B29; Fischer et al., 1998c). A cold period in
northern Greenland corresponding to the LIA is later than
reconstructed for the entire Northern Hemisphere by Mann
et al. (1998), with lowest values during AD 1620–1780 and
in the first half of the 19th century. Interestingly, the warmest
mean values of the last 1000 years at AD 1420 lie within the
time frame of the LIA.
A distinct but, compared to other periods, not exceptional
warm event in the early 20th century corresponds to the
ETCW. Since the 1870s, the values are above the 1000-year
mean. At the end of the 20th century, the temperature stagnates at a high mean level. However, as the NGT cores were
drilled between AD 1993 and 1995, the warmest years of the
recent decades (Wood et al., 2010) are not included in our
record.
For the NG stack as well as most of the individual NGT
cores, the isotopically warmest periods besides the AD 1420
event were in the 10th and 20th centuries, in particular between AD 1900 and 1950. These years are even warmer than
the most recent years covered by the NGT cores (i.e., the
1980s).
To place the results in an Arctic-wide context, we compare our northern Greenland temperature record (NG stack)
to ice core records from the Russian Arctic (Akademii Nauk
Clim. Past, 12, 171–188, 2016
– AN; Opel et al., 2013), Canada (Agassiz Ice Cap – Agassiz;
Vinther et al., 2008), Svalbard (Lomonosovfonna – Lomo;
Divine et al., 2011) and southern Greenland (Dye3; Vinther
et al., 2006b), as well as a multi-proxy reconstruction of
annual Arctic SAT (Arctic2k; Pages 2k Consortium, 2013;
Fig. 6) that covers our time period.
Note that some of these time series (Agassiz, Arctic2k)
are also stacked records with a wider regional representativeness, whereas others are single records (Dye3, AN, Lomo),
which influences the strength of correlation due to different
signal-to-noise ratios. For the discussion of the temperature
record, we concentrate on the smoothed values (30-year running means).
The strongest correlations with our NG stack are found
for the Agassiz and Arctic2k records (r = 0.58 and 0.66, respectively). For the latter, we have to consider that some of
the NGT cores (B16, B18 and B21 on the old timescale) are
used to generate this multi-proxy record. In total, 59 records
including 16 ice cores were used. NGT cores represent only
3 out of these 59 records. The correlation coefficient between
the stacked anomalies of B16, B18 and B21 and the Arctic2k
temperature is small (r = 0.24), so we can assume that the
NGT records do not dominate the reconstruction.
We conclude that a good correlation between the NG stack
and the Arctic2k record shows that the temperature in northern Greenland generally follows the Arctic-wide mean temperature.
The Lomonosovfonna record is interpreted as a winter
record and has only a weak correlation with the NG stack
(r = 0.22). More summer snow in northern Greenland compared to Lomonosovfonna could be one possible explanation
for the weak correlation between both records. While for the
other drill sites we have comparable r values for both substacks as for the NG stack, the Lomonosovfonna record has
a stronger correlation to stack “East” (r = 0.2) than to stack
“Divide” (r = −0.12), which supports the argument of different moisture sources or seasonal distribution of snowfall
in the northeast of Greenland.
The Lomonosovfonna, Akademii Nauk and Arctic2k
records show significantly more enriched δ 18 O values during the MCA. However, smaller events of abnormal warm
temperatures during the MCA are observed for Agassiz and
Dye3. Our NG stack shows warmer values earlier than the
MCA time period given by Mann et al. (1998). We conclude
that, further north in the Arctic, the warm events during MCA
may be less pronounced or earlier in timing.
The Lomonosovfonna and Arctic2k records show a dominant cold period during the LIA from AD 1580 to 1870. Also,
our northern Greenland ice cores, as well as those of Agassiz
and Akademii Nauk, reveal distinct LIA cooling periods in
contrast to the Dye3 ice core from southern Greenland. Like
in our NG stack, the cooling appears in two phases and some
decades later than described by Mann et al. (2009). For the
NG stack, the younger phase (AD 1800–1850) is of minor
amplitude and shorter duration.
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S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
183
Figure 6. Thirty-year running mean for δ 18 O values from different Arctic regions – northern Greenland (NG stack, this study), southern
Greenland (Dye3; Vinther et al., 2006b), Canada (Agassiz Ice Cap, Agassiz; Vinther et al., 2008), Russian Arctic (Akademii Nauk, AN; Opel
et al., 2013) and Svalbard (Lomonosovfonna, Lomo, Divine et al., 2011) – as well as a reconstructed record (Arctic2k; Pages 2k Consortium,
2013). All records are given on z-level scales (centered and normalized data). The correlation coefficient for the smoothed values with regard
to our stack is also given.
Between AD 1920 and 1940, there was a major warming
period in the Arctic, known as the ETCW, which was observed in all records shown here. Chylek et al. (2006) determined from meteorological data that the 1920–1930 warming was stronger than the 1995–2005 warming. For the NG
stack and Akademii Nauk record, the ETCW was warmer
than the second half of the 20th century, which distinguishes
them from other shown records. The ETCW is assumed to
be independent of external forcing but caused by internal
climate variability, in particular sea-ice–atmosphere feedbacks (Wood and Overland, 2010). This led us to conclude
that northern Greenland may also be a good place to study
forcing-independent (i.e., internal) climate changes.
However, natural external forcing (i.e., insolation, solar irradiance and volcanic eruptions) is assumed to influence the
temperature that can be studied from northern Greenland’s
ice cores.
In general, higher solar activity causes higher temperatures
(as during the MCA), whereas cold periods (e.g., LIA) are
dominated by lower solar activity (Ammann et al., 2007).
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Based on some of the NGT records (B16, B18, B21 and
B29), Fischer et al. (1998c) explained most of the long-term
variation in northern Greenland through changes in solar activity.
Volcanism causes strong negative radiative forcing
(Robock, 2000). It is assumed that volcanic eruptions inject large quantities of sulfur-rich gases into the stratosphere
and global climate can be cooled by 0.2–0.3 ◦ C for several
years after the eruption (Zielinski, 2000). Results from Crowley (2000) indicate that volcanism generally explains roughly
15–30 % of the variability in global temperatures.
Miller et al. (2012) argued that century-scale cold summer
anomalies, of which the LIA represents the coldest, occur
because natural forcing is either weak or, in the case of volcanism, short-lived. Pages 2k Consortium (2013) shows that
periods with strong volcanic activity correspond to a reduced
mean temperature. The LIA may be therefore caused by a
50-year-long episode of volcanism and kept persistently cold
because of ocean feedback and a summer insolation minimum.
Clim. Past, 12, 171–188, 2016
184
S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
Figure 7. The northern Greenland stack (NG stack; blue: annual;
dark blue: smoothed) is shown with possible forcing factors. Green
represents the reconstructed total solar irradiance (Steinhilber et al.,
2009), purple the reconstructed August Arctic sea-ice extent (Kinnard et al., 2011), and red at the bottom the stratospheric sulfate
aerosol injection for the Northern Hemisphere (Gao et al., 2008).
All values are 40-year low-pass-filtered. The discussed AD 1420
event is marked in beige.
Between about AD 1100 and 1600 we observe quasiperiodic (60–80 a) cold and warm anomalies in the NG stack
which are not present in the other Arctic records shown
(Fig. 6). The main period determined using Fourier decomposition between AD 1100 and 1600 for 30-year running
mean smoothed values is calculated with 76.31 a.
The Atlantic Multidecadal Oscillation (AMO) could be
one possible influence causing these low-frequency oscillations. Chylek et al. (2012) explain that the AMO is visible
in δ 18 O values from central Greenland. As the AMO index
reconstruction (Gray et al., 2004) does not cover the time between AD 1100 and 1600, we can only speculate about an influence in that time due to the similar periodicity. For the time
period AD 1567–1990, the correlation between the NG stack
and the AMO index is weak (r = 0.06), which might be due
to the uncertainties in historical AMO data. However, after
AD 1800 we observe a higher correlation coefficient (r = 0.
66, p < 0.05) implying a possible relation.
One of these warmer periods is at about AD 1420 ± 20,
an abnormal warm event which is observed in our northern
Greenland record and has not been pointed out in other ice
core studies before. The event is observable in all nine NGT
cores covering this time (Fig. 2) as well as in NGRIP but not
in the isotope records from southern Greenland such as the
Dye3 ice core (Fig. 6). One reason here might be the specific
geographical position in the north.
Furthermore, we observe a difference between the Canadian and Russian Arctic regarding the 1420 event. Unlike
the Russian Akademii Nauk ice core, the δ 18 O values of the
Agassiz cores from Ellesmere Island also show a tendency
towards more enriched values in that period but which is not
as strong as in northern Greenland.
Clim. Past, 12, 171–188, 2016
The fact that the 1420 event is not clearly noticeable in
other surrounding Arctic ice cores emphasizes that this event
may have occurred on a smaller regional scale. However, it
seems to have been of dominant influence and is also reflected in a smaller warming for the Arctic2k record (Fig. 6).
The spatial distribution of the 1420 event in northern
Greenland is mapped in Fig. 3b. The event is strongest in
the upper north and shows a different pattern than the δ 18 O
anomalies of the 1920/1930 warm phase, which is also attributed to internal variability and is strongest in the northeast
of Greenland.
Figure 7 shows possible forcing factors that might be related to the AD 1420 event. According to the reconstructed
total solar irradiance record of Steinhilber et al. (2009), there
was no solar maximum observed for AD 1420 that could explain the warmer temperatures in northern Greenland. As we
see no forcing anomaly, we interpret the 1420 event as likely
be caused by internal Arctic climate dynamics with a seaice–atmosphere feedback.
Box (2002) argued that climate variability in Greenland
is linked to the North Atlantic Oscillation (NAO), volcanism and sea-ice extent. NAO (Vinther et al., 2003) is calculated to be weakly reflected (r = −0.2, p < 0.01) in the NG
stack, similar to the results of White et al. (1997) for summit
ice cores, whereas none of the single NGT records is significantly correlated (p < 0.05) with the NAO index. The NG
stack has an increased signal-to-noise level, which is why
the correlation here might be clearer than from individual
records. Also, the sub-stacks of the records on the ice divide
(stack “Divide”) as well as those east (stack “East”) are significantly correlated (r = −0.19 and −0.17, p < 0.05) with
the NAO index. The cores east of the main ice divide are expected to be out of the major cyclonic track. We conclude
that NAO is not of major importance for northern Greenland
δ 18 O values.
Around AD 1420, an anti-correlation between sea-ice extent in the Arctic Ocean (Kinnard et al., 2011) and the δ 18 O
values is observed (Fig. 7). The sea-ice extent reconstruction of Kinnard et al. (2011) is based on 69 proxy records, of
which 22 are δ 18 O records. Out of these 22 δ 18 O, records 5
(NGRIP, B16, B18, B21 and B26) are also used in our NG
stack. We do not expect circular reasoning in the interpretation of the 1420 event because B16 and B26 do not reach the
age of AD 1420 and we do not see a strong anti-correlation
during any other time period.
The sea ice in the Arctic Ocean shows a recession only
during that warm period in northern Greenland. A diminished sea-ice extent would cause higher temperatures on a
regional scale and would increase the amount of water vapor
from local sources. Therefore, compared to distant sources,
more isotopically enriched moisture (Sime et al., 2013) may
contribute to precipitation in northern Greenland, in particular east of the main ice divide.
However, we do not see any direct relationship between
sea-ice extent and our NG stack during the rest of time, which
www.clim-past.net/12/171/2016/
S. Weißbach et al.: Spatial and temporal oxygen isotope variability in northern Greenland
does not exclude the relationship between sea-ice extent and
δ 18 O in northern Greenland. The sea-ice reconstruction used
is Arctic-wide, which means that the climatic events of regional extent, like an additional moisture source for northern
Greenland’s δ 18 O, do not always have to be reflected. Nevertheless, the recent NEEM δ 18 O record from northwestern
Greenland also shows a generally close relationship with the
Labrador Sea/Baffin Bay sea-ice extent (Masson-Delmotte et
al., 2015; Steen-Larsen et al., 2011).
4
Conclusions
With the full set of the NGT records, it was, for the first time,
possible to describe regional differences in the δ 18 O values
in northern Greenland over the last 1000 years.
Because of the ice sheet topography we see a clear eastto-west difference in northern Greenland δ 18 O distribution.
In total, 12 % of the spatial δ 18 O variability is attributed to
ice sheet topography. The east-to-west gradient is larger than
the north-to-south gradient. We find a more pronounced persistence of warm or cold events east of the main ice divide
and assume more stable climate conditions there. The eastern part is more influenced by local effects like changes in
the Arctic Ocean. However, these findings need to be supported by the results of climate models. For the first time, a
local warm event at AD 1420 ± 20 has been pointed out. We
assume an atmosphere–sea-ice feedback to be one possible
reason for this event.
Due to the shadowing effect of the main ice divide we find
the lowest accumulation rates in the northeast, whereas the
lowest mean δ 18 O values are found east of the main ice divide
north of the summit. The lowest δ 18 O mean values seem to
be independent of accumulation rate.
We have presented a new 1000-year stacked δ 18 O record
for northern Greenland covering 10 % of the area of Greenland. We found this NG stack to be representative of the
northern Greenland temperature.
Northern Greenland δ 18 O represents known climatic variations of the last millennium. We see a warm MCA and can
derive distinct LIA cooling from our NG stack.
The results of single-site ice core studies are likely weakened by the finding that there is only 22 % common variability in the 13 NGT cores.
The solar activity and internal Arctic climate dynamics are
likely the main factors influencing the temperature in northern Greenland. In contrast, we could not find a general cooling effect of volcanic eruptions in our data.
Acknowledgements. Stefanie Weißbach was financed by the
“Earth System Science Research School (ESSReS)”, an initiative
of the Helmholtz Association of German Research Centres (HGF)
at the Alfred Wegener Institute (AWI), Helmholtz Centre for Polar
and Marine Research.
Anna Wegner acknowledges REKLIM for funding.
www.clim-past.net/12/171/2016/
185
This study contributes to the Eurasian Arctic Ice 4k project funded
by the Deutsche Forschungsgemeinschaft (grant OP 217/2-1
awarded to Thomas Opel).
Many thanks to the drill and lab team who have measured the
δ 18 O (76.930 samples) over more than 20 years. We also thank
Johannes Freitag and Katja Instenberg for high-resolution (CT)
density measurements; Martin Rückamp for compiling the maps;
Martin Werner, Thomas Laepple and Johannes Freitag for helpful
discussions that improved the manuscript; and Kirstin Meyer for
checking the language.
We acknowledge the constructive comments from the two anonymous referees and the editor.
The NG δ 18 O stack as well as the δ 18 O and accumulation records
of all individual NGT ice cores used in the paper is available at
doi:10.1594/PANGAEA.849161.
Edited by: V. Masson-Delmotte
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