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Phil. Trans. R. Soc. A (2009) 367, 21–48
doi:10.1098/rsta.2008.0203
Published online 16 October 2008
REVIEW
Pliocene palaeoceanography of the Arctic
Ocean and subarctic seas
B Y J ENS M ATTHIESSEN 1, * , J OCHEN K NIES 2 , C HRISTOPH V OGT 3
1
AND R UEDIGER S TEIN
1
Alfred Wegener Institute for Polar and Marine Research (AWI),
27515 Bremerhaven, Germany
2
Geological Survey of Norway (NGU), 7491 Trondheim, Norway
3
Crystallography, Department of Geosciences, University of Bremen,
28359 Bremen, Germany
The Pliocene is important in the geological evolution of the high northern latitudes. It
marks the transition from restricted local- to extensive regional-scale glaciations on the
circum-Arctic continents between 3.6 and 2.4 Ma. Since the Arctic Ocean is an almost
land-locked basin, tectonic activity and sea-level fluctuations controlled the geometry of
ocean gateways and continental drainage systems, and exerted a major influence on the
formation of continental ice sheets, the distribution of river run-off, and the circulation
and water mass characteristics in the Arctic Ocean. The effect of a water mass exchange
restricted to the Bering and Fram Straits on the oceanography is unknown, but
modelling experiments suggest that this must have influenced the Atlantic meridional
overturning circulation. Cold conditions associated with perennial sea-ice cover might
have prevailed in the central Arctic Ocean throughout the Pliocene, whereas colder
periods alternated with warmer seasonally ice-free periods in the marginal areas. The
most pronounced oceanographic change occurred in the Mid-Pliocene when the
circulation through the Bering Strait reversed and low-salinity waters increasingly
flowed from the North Pacific into the Arctic Ocean. The excess freshwater supply might
have facilitated sea-ice formation and contributed to a decrease in the Atlantic
overturning circulation.
Keywords: Arctic Ocean; Pliocene; palaeoceanography; glaciation history;
tectonic activity; climate optima
1. Introduction
The Pliocene in the high northern latitudes was a time of considerable change
in climate and environment. Superimposed on the long-term cooling since
the Mid-Miocene (e.g. White et al. 1997, 1999; Bezrukova et al. 1999; Demske
et al. 2000), distinct excursions to warmer climates structured this relatively
*Author for correspondence ([email protected]).
One contribution of 11 to a Theme Issue ‘The Pliocene. A vision of Earth in the late twenty-first
century?’.
21
This journal is q 2008 The Royal Society
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22
J. Matthiessen et al.
180°
BES
Alaska
1
2
EES
BMB
Canada
CAA
3
90° W
AB
4
BB
Lo
mo
no
Tr
an
so
vR
sp
ola
idg
rD
e
rif
t
Beaufort
Gyre
LAS
90° E
KS
Siberia
5
BS
Greenland
EG
C
LS
6
8
9
NS
GSR
7
FS
Scandinavia
NC
10
GSR
0°
Bathymetric and topographic tints
–5000
–1000
–100
–10 0
100
500
1000 (m)
Figure 1. Bathymetry of the Arctic Ocean ( Jakobsson et al. 2000) and major surface currents
(black: warm currents; white: cold currents). The summer (thin white line) and winter (thick white
line) extent of sea-ice cover is shown (AB, Arctic Basin; BB, Baffin Bay; BS, Barents Sea; BES,
Bering Strait; BMB, Beaufort–Mackenzie Basin; CAA, Canadian Arctic Archipelago; ESS, East
Siberian Sea; FS, Fram Strait; GSR, Greenland–Scotland Ridge; KS, Kara Sea; LS, Labrador
Sea; LAS, Laptev Sea; NS, Nordic Seas; NC, Norwegian Current; EGC, East Greenland Current).
The locations of shallow-marine sediments that partly record warmer water conditions and ODP
holes are shown (1, North Alaska Coastal Plain, Brigham-Grette & Carter (1992) and Cronin et al.
(1993); 2, Beaufort–Mackenzie Basin, McNeil et al. (2001); 3, Meighen Island, McNeil
(1990); 4, Hvitland Beds, Fyles et al. (1998); 5, Kap København, Funder et al. (2001) and Cronin
et al. (1993); 6, ODP Hole 911A; 7, ODP Hole 909C; 8, Île de France, Bennike et al. (2002); 9,
Lodin Elv, Feyling-Hanssen et al. (1983) and Funder (1989); 10, Baffin Island (Clyde Foreland,
Qivituq Peninsula), Feyling-Hanssen (1985) and Cronin et al. (1993)).
short geological epoch in the Arctic region (Matthews & Ovenden 1990;
Brigham-Grette & Carter 1992; Cronin et al. 1993; Funder et al. 2001; Elias et al.
2006). A relatively warm Early Pliocene with small-scale glaciations was
succeeded by a progressive intensification in the Mid-Pliocene, leading to the
build-up of large ice sheets in North America and Eurasia (e.g. Fronval & Jansen
1996; Mudelsee & Raymo 2005). Since the Arctic Ocean was in the Pliocene, as it
is today, an almost land-locked basin with mostly small and shallow conduits to
the World Ocean except for the deep Fram Strait that connects the eastern
Arctic Ocean and the North Atlantic Ocean via the Nordic Seas (figure 1),
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Review. Pliocene Arctic Ocean
23
changes in the geometry of the Arctic gateways by sea-level fluctuations and
tectonic activity must have exerted a strong influence on the palaeoenvironmental evolution.
Furthermore, the hydrological cycle of the Arctic region is characterized by a
large input of freshwater to the Arctic Ocean by river run-off and import through
the Bering Strait, which, together with net precipitation, lead to a low-salinity
surface water layer (Serreze et al. 2006). Heat loss at the sea surface results in the
formation and maintenance of the sea-ice cover. The formation of a halocline
below the mixed layer insulates the saline and warmer intermediate waters from
the atmosphere and thus supports the sea-ice cover (e.g. Rudels et al. 1996).
Freshwater is exported through the gateways and may influence the Atlantic
meridional overturning circulation (AMOC) by altering deep convection in the
North Atlantic Ocean (Dickson et al. 2007).
Despite the undisputed role of the Arctic Ocean and subarctic seas in the
modern and Pliocene climate system (Driscoll & Haug 1998; Haug et al. 2001,
2005), the Arctic has attracted much attention only in the past few years since
the public has become aware of the ongoing fundamental change in the Arctic
cryosphere as a possible response to global warming, which will have a major
impact on ecosystems and human society (e.g. IPCC 2007). A number of Ocean
Drilling Program (ODP) holes have been drilled in the marginal Arctic Ocean
and subarctic seas in the past 25 years, recently complemented by the first
scientific drill holes in the central Arctic Ocean (IODP Expedition 302, ACEX—
‘Arctic Coring Expedition’, Backman et al. 2006), but these datasets have rarely
been summarized to comprehensively reconstruct the Pliocene palaeoclimate and
palaeoceanographic evolution of the high northern latitudes (Mudie et al. 1990).
Therefore, we attempt to give an overview on the current knowledge of Pliocene
palaeoceanography of the Arctic Ocean and subarctic seas as well as the major
external forcing factors (tectonic activity, freshwater supply) that may have
influenced water mass characteristics, circulation and sea-ice cover.
2. Pliocene chronostratigraphy
The chronostratigraphy of Pliocene sediments in the high northern latitudes is
principally based on magnetostratigraphy supported by age datums of various
microfossil groups. Stable isotope records on benthic and planktic foraminifers
play only a subordinate role owing to strongly discontinuous records (Fronval &
Jansen 1996; Haug et al. 2005). In the Atlantic sector of the high northern
latitudes, calcareous nanofossils and planktonic foraminifers and to a small
extent diatoms and radiolarians provide the basic age control, whereas in the
North Pacific Ocean biosiliceous microfossils are more important. Stratigraphic
resolution of calcareous and biosiliceous microfossils decreases strongly with
increasing latitude, whereas palynomorphs have a larger although not fully
explored stratigraphic potential in the Arctic Ocean (e.g. Eldholm et al. 1989;
Srivastava et al. 1989; Rea et al. 1995; Thiede et al. 1996; Raymo et al. 1999;
Backman et al. 2006).
The chronostratigraphy of central Arctic Ocean (CAO) sediments has been
intensively debated for more than 30 years (see Backman et al. (2004) for a
review). Pliocene sediments could not unequivocally be identified until the IODP
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24
J. Matthiessen et al.
Expedition 302 drilled on the Lomonosov Ridge in the CAO (Backman et al.
2006, 2008). This new chronostratigraphy confirms the conclusion by Backman
et al. (2004) that average sedimentation rates in the CAO were an order of
magnitude higher (cm instead of mm kaK1) than previously assumed in many
studies. Therefore, the Pliocene age of sediments that has been inferred from
millimetre-scale sedimentation rates in sediment cores obtained prior to IODP
Exp. 302 must be revised, requiring also a re-evaluation of palaeoceanographic
interpretations (e.g. Aksu & Mudie 1985; Herman et al. 1989; Scott et al. 1989;
Mudie et al. 1990; Clark 1996; Spielhagen et al. 1997; Grantz et al. 2001).
Therefore, the reconstruction of the Pliocene palaeoceanographic history of the
CAO is based on a single site that has been dated by 10Be stratigraphy and a few
biostratigraphic data (Backman et al. 2008; Frank et al. 2008).
Shallow-marine sediments of Pliocene age that are exposed on the circumArctic continents provide a wealth of palaeoenvironmental information, but
absolute age assignments are as critical as in the CAO (e.g. Brigham-Grette &
Carter 1992; Cronin et al. 1993; Harrison et al. 1999; Funder et al. 2001; McNeil
et al. 2001; Polyakova 2001; Bennike et al. 2002; figure 1). These sequences
comprise rather short periods, and radiometric and/or isotopic ages that are
required to place these records into an absolute chronology are rare.
Biostratigraphic datums are rarely well calibrated as, for example, illustrated
by the stratigraphic range of the benthic foraminifer Cibicides grossus that is
widespread during sea-level highstands in the Arctic region (Fyles et al. 1998;
McNeil et al. 2001). This species has been used as a stratigraphic marker for the
Mid- to early Late Pliocene (2.4–3.5 Ma) in a number of shallow-water sequences
(McNeil et al. 2001), but it ranges into the Early Pleistocene in deeper-water
sediments in the North Sea (King 1989) and on the Yermak Plateau (Osterman
1996). A younger last occurrence in deeper-water sediments is supported by new
87
Sr/86Sr data from the Beaufort–Mackenzie Basin, giving a stratigraphic range
of 2–1 Ma (McNeil et al. 2001). Despite chronological uncertainties, some
sequences may have been formed at approximately the same time (Matthews &
Ovenden 1990; McNeil 1990; Brigham-Grette & Carter 1992; Cronin et al. 1993;
Fyles et al. 1998; Funder et al. 2001; Elias et al. 2006; figure 2).
In this review, the stratigraphic concept of Gradstein et al. (2004) is used but a
proposed revision of the stratigraphic status of the Quaternary may lead to a
much shorter Pliocene epoch (http://www.stratigraphy.org).
3. Palaeoenvironmental boundary conditions in the Arctic region
The exchange of water masses between the Arctic Ocean and the World Ocean is
strongly controlled by the geometry of the conduits, and, in particular, tectonic
activity and fluctuating sea level may have altered transport rates of shallowand deep-water masses between the different basins. Furthermore, tectonic uplift
may have influenced the geometry of drainage basins and thus the distribution of
river run-off and the build-up of large continental ice sheets that supplied
sediments to the continental margins. Since these boundary conditions played a
more important role for the palaeoclimate and palaeoceanography in the Arctic
region than for any other ocean basin in the Pliocene, they are described in the
following sections.
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25
polarity
planktonic
foraminifers
calcareous
nanofossils
5
benthic d 18 O (‰)
4
3
palaeoenvironmental
events
modern
sea level
NN
21
4
Early
NN
19
Late
N22
NN
18
Middle
1
Pleistocene
0
Middle
epoch
Review. Pliocene Arctic Ocean
N
20/21 NN
16
3
9
3
Pliocene
KK
BI
1
NHG
8
6
PCO
NA
HB
LE
MI
7
M2–MG2
IF
NN
15
4
Early
time (Ma)
2
5
N19
NN
14
5
2
NN
13
(Lisiecki & Raymo 2005)
Late
6
Miocene
N18
N17
NN
12
(Miller et al. 2005)
7
–150 –100 –50
0
sea level (m)
50
Figure 2. Major palaeoenvironmental events in the Pliocene of the high northern latitudes. The
fluctuations of sea level relative to present are shown. The vertical bar delimits the approximate
depth range of the modern Bering Strait. Some key tectonic and palaeoenvironmental events are
listed. The stratigraphic range of selected shallow-marine deposits indicating warmer conditions is
shown (1, formation of sedimentary wedges; 2, Bering Strait closed; 3, Canadian Arctic
Archipelago closed; 4, subaerial exposure of Barents Sea; 5, modern East Greenland Current; 6,
North Pacific halocline; 7, southward flow through Bering Strait; 8, maximum strength of Atlantic
water inflow; 9, maximum strength of formation of Northern Component Water; BI, Baffin Island;
HB, Hvitland Beds; IF, Île de France; KK, Kap København; LE, Lodin Elv; MI, Meighen Island;
NA, North Alaska Coastal Plain; NHG, Northern Hemisphere Glaciation; PCO, Pliocene
Climate Optimum).
(a ) The Arctic Basin in the Pliocene
The topography of the circum-Arctic continents and the geometry of the Arctic
Ocean and its gateways to the North Pacific Ocean, the Nordic Seas, the Labrador
Sea and Baffin Bay have been changed frequently by the complex interaction
between tectonic activity and sea-level fluctuations in the Pliocene. The Arctic
Basin is particularly sensitive to sea-level fluctuations owing to the extensive shelf
seas (mostly shallower than 200 m) that make up more than 50 per cent of the
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J. Matthiessen et al.
modern Arctic Ocean (Jakobsson et al. 2003). A falling sea level will expose huge
shelf areas in the Siberian Arctic and will close the shallow gateways through the
Canadian Arctic Archipelago and the Bering Strait. Eustatic sea level varied on
average by G20 m around the modern level until 3.3 Ma, when a trend to
consistently lower maxima and minima commenced (figure 2; Miller et al. 2005).
The deep Arctic Ocean and adjacent seas have experienced little change in their
geometry since the Mid-Miocene except for a minor expansion and deepening of the
Fram Strait and the continuous sea-floor spreading in the Eurasian Basin and
the Nordic Seas (Kristoffersen 1990; Lawver et al. 1990; Jakobsson et al. 2007). The
history of the shallow-water connections to the Atlantic and Pacific Oceans and the
development of their geometry are less known (figure 2). The Bering Strait possibly
opened at the Miocene–Pliocene transition ca 5.5–5.4 Ma (Gladenkov et al. 2002;
Gladenkov 2006), although there is evidence from diatom analysis for intermittently open connections since the Early Miocene (Polyakova 2001). Flooding of
the Bering Strait was probably facilitated by large-scale tectonic events and/or
eustatic sea-level rise (Marincovich 2000). The Barents shelf was possibly
subaerially exposed in the Pliocene (Torsvik et al. 2002) and major parts probably
subsided below sea level ca 1 Ma (Butt et al. 2002). The sill depth of the Greenland–
Scotland Ridge has changed with time due to a variable activity of the Iceland
plume (Poore et al. 2006). Shallow conduits through the Canadian Arctic
Archipelago might not have existed in the Pliocene. Harrison et al. (1999) assumed
that Mid- to Upper Pliocene deposits in the Canadian Arctic form a single
depositional sequence. A drainage system that cut down through the regionally
extensive Pliocene sediment cover might have developed in the Early Pleistocene.
Fluvio-deltaic to deltaic–shallow-marine depositional environments prevailed along
the Canadian Arctic continental margin (Dixon et al. 1992; Torsvik et al. 2002).
Therefore, the Bering and Fram Straits might have been the only Pliocene
gateways to the Arctic Ocean.
Both tectonic activity and glacial erosion shaped the topography of the
circum-Arctic continents and influenced the atmospheric circulation in the
Pliocene (e.g. White et al. 1997; Zhisheng et al. 2001; Duk-Rodkin et al. 2004;
Praeg et al. 2005; Stoker et al. 2005). Tectonic and glacial activity, sea-level
change and bottom currents caused substantial hiatuses at the continental
margins (e.g. Bohrmann et al. 1990; Wolf & Thiede 1991; Brigham-Grette &
Carter 1992; Solheim et al. 1998; McNeil et al. 2001; Geissler & Jokat 2004;
Dahlgren et al. 2005; Praeg et al. 2005; Stoker et al. 2005). McNeil et al. (2001)
described a regional unconformity in the Beaufort–Mackenzie area of Arctic
Canada encompassing the period between ca 5.2 and 3.5 Ma that ‘is probably
recognizable on a circum-Arctic scale and beyond’. The unconformity extends
from the continental rise to the hinterland of the Beaufort–Mackenzie Delta and
was initiated by shelf exposure during the latest Late Miocene to earliest
Pliocene eustatic sea-level lowstand, further accentuated by tectonic uplift of the
basin margin. In the Siberian Arctic, the base Pliocene unconformity in the
Laptev Sea was probably caused by global sea-level fall at the Miocene–Pliocene
transition. The transgression towards the end of the Pliocene reached the present
shoreline and led to deposition of shallow-marine sediments (Drachev et al. 1998;
Kos’ko & Trufanov 2002). In the Late Pliocene to Early Pleistocene, large coastal
regions of the Eurasian Arctic were flooded due to tectonic subsidence
(Polyakova 2001).
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Review. Pliocene Arctic Ocean
27
Japsen & Chalmers (2000) compiled evidence for tectonic uplift along the
continental margins of West and East Greenland and Norway in the Neogene but
the timing of these events is partly controversial. Along the continental margins
of the Nordic Seas, Upper Pliocene sediments usually rest unconformably on
Upper Miocene or on a thin layer of Lower Pliocene sediments probably due to
erosion of Lower Pliocene sequences during an extensive global sea-level fall (e.g.
Eidvin et al. 2000; Ryseth et al. 2003; Hjelstuen et al. 2004; Tsikalas et al. 2005).
The Intra-Pliocene unconformity forms the base of large prograding sedimentary
wedges between Rockall Trough and the Vøring Plateau (e.g. Dahlgren et al.
2005; Praeg et al. 2005; Stoker et al. 2005). Tilting from the Early Pliocene
(!4 MaG0.5 Ma) to recent resulted in a basinward progradation of shelf-slope
wedges, from uplifts along the inner margin and from offshore highs. The
extensive development of the continental slope wedges in the Late Pliocene was
controlled by tectonic and/or climate evolution (Dahlgren et al. 2005; Praeg
et al. 2005; Stoker et al. 2005).
The completeness of the stratigraphic record in the CAO cannot be assessed
with the available data. Since IODP Exp. 302 drilled on the relatively shallow
Lomonosov Ridge in water depths of less than 1300 m and sedimentation rates
were relatively low throughout the Neogene sequence (Backman et al. 2006,
2008), Moore et al. (2006) expected to find other hiatuses than the large Eocene–
Miocene hiatus. Subsequently, Frank et al. (2008) described a short-term hiatus
in the Late Miocene. Numerous reflectors have been identified in the Pliocene
section of a synthetic seismic model, but they could not be related to hiatuses
due to the absence of distinct changes in the lithology and the present low
stratigraphic resolution (Backman et al. 2008), suggesting a rather continuous
sedimentation in the Pliocene. However, local sea-floor erosion by grounding
icebergs has removed Pliocene and Pleistocene sediments on the Lomonosov
Ridge (Frank et al. 2008).
(b ) Glaciation history of the circum-Arctic continents
Large-scale glaciations probably had an asymmetric geographical distribution
in the circum-Arctic regions in the Pliocene similar to the Late Pleistocene
because Eastern Siberia has only been covered by locally restricted mountain
glaciations rather than by large ice sheets (e.g. Prokopenko et al. 2001;
Glushkova & Smirnov 2007).
The Miocene–Pliocene transition is marked by the onset of small-scale glaciations in the Northern Hemisphere at ca 7.2–6.0 Ma that persisted until ca 3 Ma
(Fronval & Jansen 1996). St. John & Krissek (2002) assumed that glaciations on
Greenland started ca 7.3 Ma and glaciers were large enough to reach sea level in
southeastern and western Greenland (see also Solheim et al. 1998). In Baffin Bay
and the Labrador Sea, sporadic ice rafting has been dated at 8–7 Ma (Korstgård &
Nielsen 1989; Wolf & Thiede 1991). In ODP Hole 909C from the Fram
Strait an increase in the accumulation (figure 3) and a change in the composition of the coarse fraction between 7.5 and 6.9 Ma is interpreted as indicating a
strengthening of the Northern Hemisphere glaciations (NHGs; Winkler et al. 2002).
Glaciations were probably not widespread on the circum-Arctic continents in
the Early Pliocene. Geirsdóttir (2004) stated that locally restricted glacial
deposits are as old as 5 Ma but local ice caps may have already existed in the
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28
smectite/(illite+chlorite) ratio
6
0
5
4
5
3
LMG
NHG
2
10
1
0
acc. rate(>63 µm) (g cm–2 kyr–1)
J. Matthiessen et al.
15
2
6
4
8
age (Ma)
Figure 3. Ratio of the clay mineral groups smectite versus illite and chlorite (lower curve) and
accumulation rates (acc. rate) of coarse fraction (upper curve) in Hole 909C from the Fram Strait
in the Late Miocene and the Pliocene ( Winkler et al. 2002). The shaded areas indicate the NHG
and the intensification of small-scale glaciations in the Late Miocene (LMG).
Miocene. Debris flow deposits and glacial diamictons in ODP Holes 918 and 987
off east Greenland suggest a larger-scale glacial advance across the Greenland
shelf in the Early Pliocene (Solheim et al. 1998). In southeastern Alaska,
tidewater glaciers were probably present since the Late Miocene but definitely
occurred since 4.2–4.3 Ma (Krissek 1995; Rea & Snoeckx 1995; Lagoe & Zellers
1996; Duk-Rodkin et al. 2004). A Cordilleran ice sheet covered southeastern
Alaska and small local ice caps developed in northwestern Canada (Barendregt &
Duk-Rodkin 2004). In the North Pacific Ocean, ice-rafted debris (IRD) records
suggest minor glaciations in both Alaska and Kurile/Kamchatka during the
Early and Mid-Pliocene (Krissek 1995).
Both terrestrial and marine records provide consistent evidence for the
intensification of continental glaciations in the Northern Hemisphere since the
Mid-Pliocene. This transition was probably not abrupt but rather gradual
between 3.6 and 2.4 Ma and included an excursion to the Pliocene climate
optimum (Mudelsee & Raymo 2005; figure 2). In this period the continental ice
volume was probably strongly reduced (C25 m sea level) and continental ice was
absent in the Northern Hemisphere except for a small ice cap on Greenland
(Dowsett 2007). Mudelsee & Raymo (2005) suggested that slow tectonic forcing
such as closing of ocean gateways and mountain building processes were the root
cause of the long-term intensification. Although the exact timing and succession
of glaciations in the high northern latitudes is still an open issue due to
chronostratigraphic uncertainties and a low temporal resolution of most
terrestrial sequences, it is likely that there were regional differences in the size
and extent of ice sheets. In general, the first occurrence of major continental
glaciations in the subpolar regions was much later during the Mid-Pleistocene
transition (Ehlers & Gibbard 2007), pointing also to a pronounced long-term
temporal and spatial heterogeneity.
Nevertheless, the available data suggest a somewhat coeval development of
major ice sheets around the Arctic Basin since the late Mid-Pliocene.
A compilation of North Atlantic Ocean and Nordic Seas IRD records by
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Review. Pliocene Arctic Ocean
29
Kleiven et al. (2002) shows temporal and spatial differences in initial supply from
the circum-North Atlantic ice sheets since 3.6 Ma, but fluctuations of all major
ice sheets were synchronous since 2.75–2.72 Ma. This coincides with the onset of
large-scale IRD deposition in the North Pacific Ocean at 2.73 Ma (Krissek 1995;
Haug et al. 2005). Geirsdóttir (2004) compiled evidence for the initial development
of a major ice sheet on Iceland at ca 2.9 Ma, and subsequent intensifications at ca
2.7 Ma and from 2.4 to 2.5 Ma. The Lodin Elv and Kap København formations
contain the earliest albeit circumstantial evidence for glaciations in east Greenland
at ca 2.4–2.5 Ma (Funder 1989; Funder et al. 2001). Marine IRD records suggest an
intensification of glaciations on eastern and southern Greenland slightly earlier at
ca 3 Ma (Wolf & Thiede 1991; Kleiven et al. 2002; St. John & Krissek 2002) that
may have resulted in an ice-sheet advance onto the East Greenland Shelf (Vanneste
et al. 1995). The size of the Greenland ice sheet might have fluctuated considerably,
being reduced to small mountainous ice caps in southeast Greenland during
interglacials (Funder et al. 2001).
There is only marine evidence for an intensification of glaciations on
Scandinavia and in the Barents Sea at 2.8–2.5 Ma, but these were probably of
intermediate size until the Mid-Pleistocene (e.g. Mangerud et al. 1996; Solheim
et al. 1998). The first ice advances to the coast and onto the shelf occurred locally
since ca 2.8–2.7 Ma (Fronval & Jansen 1996; Eidvin et al. 2000; Knies et al. 2002;
Geissler & Jokat 2004). The centre of glaciations was possibly located farther
north and maximum glacial erosion occurred in the Barents Sea in the Pliocene
(Mangerud et al. 1996). The intensification of glaciations on Fennoscandia and
the Northern Barents Sea is marked by lithological changes in marine records
and increased accumulation of terrigenous sediments (Henrich et al. 1989;
Wolf & Thiede 1991; Knies et al. 2002; Hjelstuen et al. 2004).
Thus, relatively high sand contents in Hole 911A reflect the advance of the
northern Barents Sea ice sheet onto the outer shelf since 2.7 Ma (figure 4), coeval
with a substantial increase in large IRD (Knies et al. 2002). The provenance of
sediments abruptly shifted from a distal to a proximal source at 2.4 Ma because
high magnetic susceptibility and clay mineral group ratio S/(ICC) values are
associated with a particular source region in the hinterland of the Kara Sea (Stein
et al. 2004). The clay mineral group ratio indicates that sediment composition
changed from a distal Westsiberian (smectite) to a proximal Spitsbergen
provenance (illite, chlorite) at Site 911 on Yermak Plateau. A comparable trend
at Site 909 in the Fram Strait with a decrease of the clay mineral group ratio and an
increase of coarse fraction accumulation rates during the intensification of NHG has
been attributed to increased glacial weathering and erosion on the adjacent
continents (figure 3; Winkler et al. 2002). This implies that sediment supply from
the adjacent ice sheets on northeastern Greenland and in the northwestern Barents
Sea almost simultaneously increased since 2.3–2.4 Ma.
Duk-Rodkin et al. (2004) reconstructed regional-scale Cordilleran and local ice
caps in northern Alaska and the Beaufort–Mackenzie area between 2.6 and
2.9 Ma. IRD deposition started in the North Pacific at 2.73 Ma (e.g. Haug et al.
2005) and the provenance of IRD points to SE Alaska as one major source area
(McKelvey et al. 1995). The Atlanta till in the mid-USA (2.41G0.14 Ma) is
possibly the oldest direct evidence of continental glaciations in the Northern
Hemisphere and was nearly the most southerly advance of the Laurentide Ice
Sheet (Balco et al. 2005).
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J. Matthiessen et al.
lithologic/ magnetic
seismic
susceptibility
units
(10–5 SI)
1000
0
0.5 1.0 1.5
0
50
100
1.5
2.0
IA/YP3
Late
2.0
PS
2.5
2.5
PCO
IB/YP2
3.0
Pliocene
Middle
age (Ma)
500
dinoflagellate
cysts (%)
Pleistocene
Early
0
1.5
ratio
sm /(il+chl)
3.0
3.5
3.5
0
20
40
sand (wt%)
0
5
10
carbonate (%)
Figure 4. The Mid-Pliocene to Lower Pleistocene section of ODP Hole 911A. The shaded areas
mark the local Pliocene climate optimum and the change in provenance of sediments (PS) in the
Late Pliocene. The age model is based on magnetostratigraphy and calcareous nanofossils (Thiede
et al. 1996; Sato & Kameo 1996; Sato et al. 2004). (Sources of data: lithological units IA and IB,
Thiede et al. 1996; seismic units YP3 and YP2, Geissler & Jokat (2004); magnetic susceptibility,
Thiede et al. 1996; dinoflagellate cysts, Matthiessen & Brenner (1996) and Knies et al. (2002);
TOC, Stein & Stax (1996) and Knies et al. (2002); clay mineral group ratio smectite/(illiteC
chlorite), sm/(ilCchl), and sand contents, Knies et al. (2002)). Circles, Brigantedinium spp.;
squares, F. filifera s.l.; crosses, O. centrocarpum.
The intensification of glaciations coincided with tectonic uplift since ca 4 Ma
that, in conjunction with glacial erosion, ultimately led to the transfer of huge
amounts of sediments to the continental margins and the formation of
sedimentary wedges and deep-sea fans (e.g. Arthur et al. 1989; Krissek 1995;
Solheim et al. 1998; McNeil et al. 2001; Winkler et al. 2002; Duk-Rodkin et al.
2004; Geissler & Jokat 2004; Dahlgren et al. 2005; Praeg et al. 2005, Stoker et al.
2005). The uplifted land areas along the European continental margins might
have been the topographic precondition for nucleation and growth of glaciations
(e.g. Eyles 1996; Dahlgren et al. 2005; Tsikalas et al. 2005). The progradation and
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31
aggradation of shelves along Norway, the Barents Sea, Greenland, the Beaufort
Sea and the North Pacific Ocean since the Early Pliocene was partly related to
large-scale glaciations in these areas.
The onset of sedimentary wedge formation was not necessarily synchronous
along the high-latitude continental margins such as those off northwest Europe,
but a glacial influence on sedimentation is generally recorded since 2.74–1.6 Ma
(Dahlgren et al. 2005). The upper part of prograding wedges in the Norwegian
Sea younger than 1.5 Ma (Sejrup et al. 2005) record advances of ice sheets across
the shelves. Shelf progradation probably occurred mainly during glacial periods
in the vicinity of cross-shelf troughs where fast-flowing ice streams supplied
sediments to submarine fans. Ice streams were important transport pathways
along the Greenland, Norwegian and Arctic continental margins during the Last
Glacial Maximum, but there are few indications of earlier activity in the Pliocene
(e.g. Stokes & Clark 2001 and references therein; Geissler & Jokat 2004; Ottesen
et al. 2005, 2007; Wilken & Mienert 2006).
Furthermore, the timing and interaction of tectonic uplift and erosion of
topographic barriers influenced the availability and distribution of moisture from
different source regions that might have controlled ice-sheet growth in coastal
Alaska (Duk-Rodkin et al. 2004). Difference in timing and extent of glaciations in
northwest Canada and east-central Alaska was probably related to proximity
and size of moisture source (Duk-Rodkin et al. 2004). The renewed uplift in
coastal southeast Alaska since 4 Ma resulted in a progressive barrier to Pacific
moisture, resulting in increased continentality, and if the Arctic Ocean was ice
free, it might have been a significant moisture source for continental glaciations
(White et al. 1999; Duk-Rodkin et al. 2004).
The continuous tectonic uplift of mountain ranges in the Asian hinterland of
the Arctic Ocean, such as the Tibetan Plateau, might have influenced Arctic
climate in various ways. Zhisheng et al. (2001) compiled evidence to show that
the atmospheric circulation system was strongly changed by uplift and extension
of the Tibetan plateau in the Pliocene. Uplift of portions of the Tibetan Plateau
and ice build-up were probably the driving forces behind long-term desertification in China with distinct increases at 3.6 and 2.6 Ma (Guo et al. 2004). The
Late Pliocene uplift at Lake Baikal coincides with early glaciations in
southeastern Siberia (Prokopenko et al. 2001).
A number of Late Pliocene continental glaciations have been recorded from
northern Alaska, northern Canada, Iceland and Siberia, but extensive
glaciations might only have occurred in these regions at the Pliocene–
Pleistocene boundary ( Vanneste et al. 1995; Prokopenko et al. 2001;
Duk-Rodkin et al. 2004; Geirsdóttir 2004; Wilken & Mienert 2006). Since the
formation of reflector R7 at the Barents Sea continental margin, which marks
the intensification of NHG at 2.3–2.5 Ma, only moderate glaciations were
recorded along the Norwegian continental margin until the Early Pleistocene
(Henrich et al. 1989; Jansen & Sjøholm 1991; Fronval & Jansen 1996; Solheim
et al. 1998; Hjelstuen et al. 2007). Between ca 3 and 1 Ma a high-frequency and
low-amplitude variability of IRD deposition occurred on obliquity time scales,
suggesting reduced glacial activity (stable ice margins) around the Nordic Seas
(Bohrmann et al. 1990; Jansen et al. 2000). Along the European continental
margin from Ireland to Svalbard ice sheets reached the shelf edge earlier in the
north in the latest Pliocene and advanced to the continental slopes first since
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J. Matthiessen et al.
the Mid-Pleistocene while glacial advances were more frequent only in the past
0.5 Ma (Solheim et al. 1998; Knies et al. 2002, 2007; Geissler & Jokat 2004;
Sejrup et al. 2005).
(c ) Freshwater supply to the Arctic Ocean
In contrast to other ocean basins, freshwater supply by rivers and by inflow
through the Bering Strait together with net precipitation plays a major role in
the hydrological cycle (Serreze et al. 2006). Approximately 3200 km3 are
annually discharged by rivers into the Arctic Ocean, mainly from the Mackenzie,
Lena, Yenisey and Ob watersheds (Serreze et al. 2006). Climate change and
tectonic activity must have had a considerable impact on the geometry of
catchment areas and on the volumes of freshwater and sediments discharged into
the Arctic Ocean. Thus, Driscoll & Haug (1998) proposed that the intensification
of the North Atlantic Drift since the earliest Pliocene, due to the closure of the
Central American Isthmus, might have enhanced moisture delivery to Eurasia.
This freshwater is supplied by the Siberian rivers to the Arctic Ocean and is
further transported into the Nordic Seas and North Atlantic Ocean where it may
have a major impact on deep-water formation.
Catchment areas and river systems differed from the modern configuration
but the impact on distribution and rates of discharge cannot currently be
evaluated. Discharge rates cannot be estimated from fluvial sediment supply
because there need not be a linear relationship between fluvial supply and
freshwater charge. Today, the major Arctic rivers differ strongly in their
sediment load. Although the Mackenzie River transports the largest sediment
load of all Arctic rivers its run-off is much smaller than that of Lena, Ob and
Yenisei (Gordeev et al. 1996).
In the Barents Sea, fluvial systems might have existed during the early phase
of the sedimentary wedge growth in the Early Pliocene (Solheim et al. 1998;
Dahlgren et al. 2005). Siberia was probably drained to the Arctic Ocean in the
Pliocene (Driscoll & Haug 1998 and references therein). Franke et al. (2001)
stated that Late Miocene and Pliocene tectonic deformation associated with the
Indo–Asian collision modified the drainage pattern across the Siberian craton,
resulting in large freshwater inflow and influx of terrigenous sediment since the
early Late Miocene. Prior to the uplift of the Primorsky Range north of Lake
Baikal in the Pleistocene, the Lena drained a much larger area including Lake
Baikal whereas today it is drained by the Yenisei (Mats et al. 2000).
In the Canadian Arctic, the modern Mackenzie River system is a relatively
young feature existing since the Late Pleistocene (Duk-Rodkin & Hughes 1994).
During the Pliocene, the area drained to the Arctic Ocean was less than 10 per
cent of the Mackenzie drainage area today, suggesting that river run-off was
probably much smaller in the Pliocene. In contrast to the modern drainage
systems, the northern cordillera was the headwater of drainage systems that
reached three oceans and a much larger area of Arctic Canada was drained to the
east, in particular into the Labrador Sea (Duk-Rodkin & Hughes 1994).
Greenland might have been an additional significant freshwater source for the
Labrador Sea during interglacials when the ice sheet was absent and the island
was drained to the west (Funder et al. 2001).
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4. Pliocene palaeoceanography
The predominant siliciclastic sediments that have low biogenic carbonate, opal
and marine organic carbon contents limit our ability to reconstruct changes in
water mass characteristics, surface-water productivity and sea-ice cover of the
Arctic Ocean (e.g. Thiede et al. 1996; Stein & Stax 1996; Backman et al. 2006).
The interpretation of the records is further complicated by diagenetic processes
leading, for example, to the formation of authigenic carbonates, manganese and
iron oxides, and iron sulphides in the sediments (e.g. Thiede et al. 1996; Chow
et al. 1996; Henrich et al. 2002; Expedition 302 Scientists 2006; Haley et al.
2008a; Krylov et al. 2008; St. John 2008). Sedimentological, petrographical and
mineralogical proxies have not yet been fully explored, but recent studies
demonstrate that they might be useful for palaeoceanographic reconstructions
(e.g. Darby 2008; Krylov et al. 2008; St. John 2008).
Pliocene marine deposits have been recovered from the marginal Arctic Ocean
and from isolated outcrops in the circum-Arctic hinterland. The shallow-marine
sequences cover only short periods in the Pliocene (e.g. Yakhimovich et al. 1990;
Brigham-Grette & Carter 1992; Mullen & McNeil 1995; Thiede et al. 1996; Fyles
et al. 1998; Harrison et al. 1999; Funder et al. 2001; McNeil et al. 2001; Polyakova
2001; Bennike et al. 2002) and therefore sites drilled during ODP Legs 104, 105,
145, 151 and 162 in the subarctic seas still provide the majority of palaeoenvironmental information. Palaeoceanographic reconstructions in the high
northern latitudes are furthermore biased to the period since the intensification
of NHG in the Mid-Pliocene (e.g. Henrich et al. 2002; Haug et al. 2005). The
Early Pliocene did not receive much attention probably owing to more equally
warm climate conditions.
(a ) Large-scale circulation in the high northern latitudes
The large-scale circulation in the Pliocene was similar to that of today. The
modern Arctic estuarine circulation with a surface-water outflow and deep inflow
through the Fram Strait was probably established in the Early Miocene (Jakobsson
et al. 2007). Radioisotope and mineralogical studies suggest that the Transpolar
Drift and possibly also the Beaufort Gyre have been persistent features since the
Mid-Miocene (Darby 2008; Haley et al. 2008b; Krylov et al. 2008; St. John 2008).
Interpretations differ with respect to the contribution of the different current
systems to the circulation in the CAO, being dominated either by the Transpolar
Drift (Krylov et al. 2008; Haley et al. 2008b) or by the Beaufort Gyre circulation
(Darby 2008). In the Nordic Seas, relatively warm waters flowed northwards along
the Scandinavian continental margin whereas colder waters were exported along the
Greenland continental margin from the Arctic Ocean to the North Atlantic Ocean
(Henrich et al. 1989, 2002; Bohrmann et al. 1990; Wolf & Thiede 1991; Fronval &
Jansen 1996). Similarly, the North Atlantic Drift penetrated continuously into the
Labrador Sea in the Pliocene (de Vernal & Mudie 1989).
The most significant change in the surface circulation occurred in the MidPliocene. The circulation through the Bering Strait reversed from a southward
flow between the opening at 5.5–5.4 and 3.6 Ma to a northward flow after 3.6 Ma
as indicated by the migration of molluscs between the North Atlantic–Arctic
Oceans and the North Pacific Ocean (Marincovich & Gladenkov 2001;
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J. Matthiessen et al.
Gladenkov et al. 2002; Gladenkov 2006). The reversal in flow direction might
have been an effect of the oceanographic changes associated with the closure of
the Central American Seaway until 2.7 Ma, which may have further induced an
intensification of the North Atlantic Drift and the formation of North Atlantic
deep water (Haug et al. 2001). The reversal in flow direction is associated
with the onset of a prominent sea-level fall that might have led to the
short-term closure of the Bering Strait during marine isotope stages M2–MG2 at
3.3 Ma (figure 2).
The intermediate- and deep-water circulation was more strongly affected by
the variable sill depths of the gateways (e.g. Bohrmann et al. 1990; Henrich et al.
2002; Poore et al. 2006). Neodymium isotopes indicate that Arctic intermediate
waters originated from brine formation in the Eurasian shelf regions in the
Pliocene rather than being advected from the North Atlantic Ocean as in the
past 2 Ma (Haley et al. 2008a). The deep-water overflow from the Nordic Seas to
the North Atlantic Ocean was enhanced in the Early and Mid-Pliocene (Henrich
et al. 2002; Poore et al. 2006). This phase might have been linked to the
southward surface-water flow through the Bering Strait. Since the NHG
intensified and a northward flow through the Bering Strait was established,
water-mass exchange between the North Atlantic Ocean and the Nordic Seas
diminished (Henrich et al. 2002; Poore et al. 2006). Some deep-water formation
might have been possible by brine rejection during sea-ice formation.
Apart from the Bering Strait and the Greenland–Scotland Ridge, the possible
impact of a change in the cross-sectional area of other gateways on the
circulation is not considered in palaeoceanographic reconstructions. Since the
Barents Sea and the conduits through the Canadian Arctic Archipelago were
probably closed in the Pliocene (Harrison et al. 1999; Butt et al. 2002; Torsvik
et al. 2002), the configuration of the Arctic Basin and its gateways differed
considerably from the modern situation. The Bering and Fram Straits as well as
the Greenland–Scotland Ridge formed the major bottlenecks for the exchange
of water masses with the adjacent North Pacific and Atlantic Oceans in the
Pliocene. Sea-level change must have further led to a recurrent opening and
closing of the shallow (less than 50 m) Bering Strait, and exposure and flooding
of the large shallow Siberian shelves. Since the global sea level fluctuated by
G50 m relative to present in the Pliocene (figure 2; Miller et al. 2005), the
geometry of the Bering Strait must have ranged from a closed state to a much
larger cross-section than today if isostatic movements could be neglected.
Short-term closure might have occurred during pronounced sea-level lowstands
at 4.9, 4.0, 3.3 and 2.5 Ma (Miller et al. 2005), but sea level was often higher in
the Early Pliocene than today, enabling an increased water-mass exchange.
Since 3.6 Ma sea level steadily decreased and the Bering Strait might have been
closed more frequently.
The results of modelling experiments suggest that the closure of Arctic
gateways must have influenced the circulation and distribution of freshwater in
the Arctic and North Atlantic Oceans (e.g. Butt et al. 2002; Wadley & Bigg 2002;
Hu et al. 2007). On the basis of an Early Pleistocene geometry of the Arctic Basin
with an open Canadian Arctic Archipelago but a closed Barents Sea, Butt et al.
(2002) calculated an increased input of relatively warm Atlantic waters into the
Arctic Ocean. The closed shallow seaways through the Canadian Arctic
Archipelago prevent a southward flow of low-salinity Arctic surface waters to
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Baffin Bay and the Labrador Sea (Wadley & Bigg 2002). The AMOC will
increase due to increased deep-water formation in the Labrador Sea. At the same
time, the increased freshwater export from the Arctic Ocean through the Fram
Strait may suppress deep-water convection in the Nordic Seas. Closing of the
Bering Strait may lead to increased surface salinities in the Arctic Ocean and
adjacent Nordic Seas because low-salinity waters are transported northwards
through the open Bering Strait, which probably cause a decrease of the strength
of the AMOC (Shaffer & Bendtsen 1994; Wadley & Bigg 2002).
The response of the AMOC to freshwater perturbations is today strongly
affected by the status of the Bering Strait (Hu et al. 2007). The AMOC recovers
faster when the Bering Strait is open because this freshwater is then exported
both to the South Atlantic and via the Arctic Ocean to the North Pacific Ocean.
Freshwater discharge to the North Atlantic Ocean while the Bering Strait is
closed will lead to a recirculation in the Arctic Ocean, an enhanced export of
freshwater and sea ice and a prolonged period of recovery of the AMOC
(Hu et al. 2007). Since glacial–interglacial cycles are coupled to sea-level
fluctuations and thus either an open or closed Bering Strait, the contrasting
model scenarios might correspond to long-lasting instabilities in the meridional
circulation during glacials in contrast to a rather stable interglacial situations
(De Boer & Nof 2004; Hu et al. 2007). Freshwater fluxes need not be continuous
to prevent a recovery if the Bering Strait is closed, but an open Bering Strait
requires a continuous flux to prevent deep convection from recovery. However,
the exact sea-level history of the Bering Strait is unknown so it is not possible to
derive the potential temporal variability during glacial–interglacial cycles.
(b ) Variability of water masses in the high northern latitudes
The reconstruction of the Pliocene palaeoceanography of the Arctic Ocean and
subarctic seas is strongly hampered by a low production and a variable
preservation of planktonic and benthic microfossils. Thus, the stable oxygen and
stable carbon isotope records of planktonic and benthic foraminifers contain large
gaps in the high northern latitudes (e.g. Fronval & Jansen 1996; Haug et al. 2005).
The discontinuous benthic isotope record from ODP Sites 642/644 (Vøring
Plateau, eastern Nordic Seas) is the exception to the rule and reflects distinct
glacial excursions in the past ca 6.5 Ma. Large fluctuations indicating long-lasting
changes in either global ice volume or bottom-water temperature or both are
recorded between 6.5 and 2.5 Ma (Fronval & Jansen 1996). The gradual cooling
since the Late Miocene is also reflected in various planktonic and benthic
microfossil records (e.g. Locker & Martini 1989; Mudie et al. 1990; McNeil et al.
2001), whereas in the Labrador Sea and Baffin Bay relatively cold conditions
prevailed throughout the Pliocene (Aksu & Kaminski 1989; de Vernal & Mudie
1989). A pronounced expansion of cold surface waters is reflected by the southward
migration of warm-adapted calcareous nanofossils in the North Pacific and
Atlantic Oceans since 2.74 Ma (Sato & Kameo 1996; Sato et al. 2004).
The absence of calcareous (foraminifers, coccolithophores) and biosiliceous
(diatoms, radiolarians, silicoflagellates) microfossils and the low abundances of
organic-walled palynomorphs in the CAO prevent the detection of any
palaeoceanographic variability in the Pliocene (Backman et al. 2006). In the
marginal Arctic Ocean, opal and carbonate contents are generally low, usually
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J. Matthiessen et al.
containing negligible amounts of biosiliceous and calcareous microfossils, whereas
palynomorphs are rather consistently present (figure 4).
In the subarctic seas, a variable deposition of biogenic opal and carbonate
reflects changing hydrographic conditions (Henrich et al. 1989, 2002; Bohrmann
et al. 1990; Haug et al. 1995; Rea & Snoeckx 1995). Shifts from opal to carbonate
accumulation were attributed to major rearrangements in circulation and watermass formation in the Nordic Seas since the Mid-Miocene (Bohrmann et al.
1990). Opal deposition and reduced carbonate preservation reflect a diminished
exchange of surface and deep waters with the North Atlantic Ocean between 5.4
and 4.8 Ma, whereas increased carbonate accumulation and preservation indicate
times of enhanced advection of relatively warm surface waters until ca 3 Ma
(Henrich et al. 1989, 2002; Bohrmann et al. 1990; Wolf & Thiede 1991; Fronval &
Jansen 1996; figure 2). This stronger inflow is linked to an increased formation of
Norwegian Sea deep water and overflow to the North Atlantic Ocean (Bohrmann
et al. 1990; Henrich et al. 2002; Poore et al. 2006).
An increased opal production in the eastern Labrador Sea was caused by iceedge diatom blooms between 4 and 2 Ma, possibly related to the advection of
cold and low-salinity sea-ice-covered waters from the Arctic Ocean to the North
Atlantic Ocean with the East Greenland Current (Bohrmann et al. 1990). The
high opal accumulation in both the Nordic Seas and the North Pacific Ocean in
the earliest Pliocene was probably associated with the global biogenic bloom due
to increased nutrient supply between 4.5 and 7 Ma (Cortese et al. 2004).
Opal accumulation broke down in the North Pacific Ocean synchronously with
increased IRD deposition since 2.74 Ma due to the onset of the permanent
stratification that reduced upwelling of nutrient-rich waters, but opal production
resumed in the latest Pliocene (Haug et al. 1995, 2005; Rea & Snoeckx 1995).
This event is associated with the development of the halocline in the North
Pacific Ocean. Seasonal contrasts in sea-surface temperature increased and the
warm pool of late summer surface water might have been a potential source of
moisture to supply the growing North American ice sheet (Haug et al. 2005;
Swann et al. 2006). A significant freshening in surface-water salinities by 2–4 psu
(practical salinity units) associated with meltwater supply has helped to
maintain the stratification (Swann et al. 2006).
Carbonate preservation in the Nordic Seas was good until 3 Ma succeeded by a
time-transgressive change to poor preservation due to a strong reduction in
warm-water inflow associated with a reduced deep-water export and the
establishment of a strong temperature gradient between the North Atlantic
Ocean and the Nordic Seas (Henrich et al. 1989, 2002; Fronval & Jansen 1996).
The sediments are virtually carbonate free in the Nordic Seas and Labrador Sea
between 2.8 and 1.9 Ma, except in the easternmost Norwegian Sea that had low
but variable contents (Henrich et al. 2002). This southeast to northwest gradient
in the Nordic Seas in the Late Pliocene might have been related to a strong
temperature gradient (Fronval & Jansen 1996). The carbonate preservation was
probably poor due to (i) low production of carbonate shells in the surface water,
(ii) sluggish renewal of deep waters induced by a rather stable sea-ice cover and/
or increased freshwater supply, and/or (iii) production of carbonate corrosive
dense brines during sea-ice formation (Henrich et al. 2002). Northern Component
Water reconstructions support a steady decline of deep-water overflow to the
North Atlantic Ocean between 3 and 2 Ma (Poore et al. 2006). Excess freshwater
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that may have helped to suppress deep-water formation in the Nordic Seas might
have been increasingly supplied from the North Pacific Ocean since (i) the onset
of the modern East Greenland Current at ca 4 Ma (Bohrmann et al. 1990),
(ii) the reversal to a northward flow through the Bering Strait at 3.6 Ma
(Marincovich & Gladenkov 2001; Gladenkov et al. 2002; Gladenkov 2006), and
(iii) the substantial decrease of sea-surface salinities in the North Pacific Ocean
at ca 2.7 Ma (Swann et al. 2006). Since the latest Pliocene, the first intrusions of
the Proto-Norwegian Current into a narrow corridor in the southeastern Nordic
Seas caused a much better carbonate preservation and a resumption of deepwater formation in the Nordic Seas (Henrich et al. 2002).
The Mid- to Late Pliocene cooling in the high latitudes was obviously
punctuated by pronounced warming events in the subarctic seas and the
marginal Arctic Ocean that were at least intermittently much warmer than
today (figure 2). Harrison et al. (1999) suggested a link between global
temperature fluctuations and Arctic transgressive–regressive depositional cycles
in the past 47 Myr, with sea-level highstands being associated with temperature
maxima. This is in agreement with previous studies demonstrating that climates
with a warmer fauna and flora are related to transgressions or sea-level
highstands often either succeeded or preceded by colder climate conditions (e.g.
Funder 1989; Brigham-Grette & Carter 1992; Funder et al. 2001; McNeil et al.
2001). The terrestrially exposed warmer-water sequences may represent rather
short interglacials such as the Kap København Formation (marine isotope stage
95/96(?)) that probably had a duration of less than 20 000 years (Funder et al.
2001). A pronounced climate cyclicity in the Pliocene of the marginal Arctic
Ocean is supported by the stable oxygen isotope composition of bivalve shelves
from the Tjörnes beds in Iceland, which shows a large-scale variability of seawater temperature superimposed on a long-term cooling trend between ca 4.3
and 2.5 Ma (Buchardt & Simonarson 2003).
Uplifted shallow-marine sequences in the marginal Arctic Ocean that indicate
warmer temperatures than today with at least seasonally ice-free conditions and
establishment of boreal forest to forest tundra have been mainly reported from the
Mid- and Late Pliocene (figure 2; Feyling-Hanssen et al. 1983; McNeil 1990;
Brigham-Grette & Carter 1992; Cronin et al. 1993; Fyles et al. 1998; Funder et al.
2001; Bennike et al. 2002). The most prominent of these events occurred in the MidPliocene between ca 3.3 and 3.0 Ma when increased meridional heat transport to
the high-latitude North Pacific and North Atlantic Oceans led to the inflow of
warmer waters into the Arctic Ocean and ice-free conditions in the marginal Arctic
Ocean (Cronin et al. 1993; Knies et al. 2002; Dowsett et al. 2005). Advection of
warmer waters is reflected by increased carbonate accumulation and abundances of
subpolar planktonic foraminifers in the North Pacific Ocean and Nordic Seas
probably due to increased production in warmer surface waters and/or reduced
dissolution (Henrich et al. 1989; Bohrmann et al. 1990; Haug et al. 1995; Rea &
Snoeckx 1995; Baumann et al. 1996). In the Nordic Seas, the carbonate maximum
strongly diminishes to the north, but a maximum of the dinoflagellate cyst
Operculodinium centrocarpum at the western and northern Barents Sea margin
may indicate an increased inflow of Atlantic waters into the eastern Arctic Ocean
(e.g. Hole 911A, figure 4; Smelror 1999; Knies et al. 2002; Ryseth et al. 2003). The
timing of this event in Hole 911A is different compared with Dowsett et al. (2005),
but this may be due to the poor age control in the lower part of Hole 911A.
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J. Matthiessen et al.
Dinoflagellate cysts principally record relatively cold, stable and productive seasurface conditions with a seasonal sea-ice cover (Brigantedinium spp.) in the
marginal eastern Arctic Ocean in the Mid- to Late Pliocene. Possibly slightly
warmer conditions with a reduced sea-ice cover as indicated by abundant
Filisphaera filifera s.l. occurred between 2.1 and 1.8 Ma (figure 4).
(c ) Arctic Ocean sea-ice cover in the Pliocene
The Arctic Ocean was probably covered by sea ice in the Pliocene, but the
spatial and seasonal extent is still a controversial issue (Cronin et al. 1993, 2008;
Polyakova 2001; Dowsett 2007; Darby 2008; Haley et al. 2008b; Krylov et al.
2008; St. John 2008). Owing to sparse microfossil records, the distinction
between a perennial and seasonal sea-ice cover relies mainly on sedimentological
and mineralogical data. The interpretation of these records is further
complicated by a possible contribution of IRD from melting icebergs. St. John
(2008), who discussed the implications of coarse fraction data (more than 250
and 150–250 mm, respectively) from Lomonosov Ridge (IODP Exp. 302), stated
that the low contents throughout the Neogene (generally less than 5%) do not
allow one to unequivocally distinguish sea ice from iceberg transport. Darby
(2008) suggested that, based on the size spectra of detrital iron oxide grains in
the same record, icebergs were of minor importance for the transport of IRD.
This may be related to restricted glaciations in the source region of the
Lomonosov Ridge sediments. The coarse fraction in CAO sediments substantially increased first in the Mid- to Late Pleistocene (Spielhagen et al. 2004),
suggesting that since then icebergs more frequently passed the Lomonosov Ridge
and the eastern Arctic Ocean. At sites proximal to Pliocene ice sheets such as the
Nordic Seas, the Fram Strait and Yermak Plateau, iceberg transport was more
important during the intensification of NHG (Henrich et al. 1989; Wolf & Thiede
1991; Winkler et al. 2002; figures 3 and 4). In general, sea ice was probably the
dominant transport mode to distribute IRD in the Pliocene of the CAO, with a
minor contribution from iceberg transport.
Sedimentological, mineralogical, radiogenic and cosmogenic isotope data
suggest that a permanent sea-ice cover existed in the CAO since the Mid-Miocene
(Darby 2008; Frank et al. 2008; Haley et al. 2008a; Krylov et al. 2008; St. John
2008). Darby (2008) and Krylov et al. (2008) suggested that, based on the
provenance of the detrital fraction of sediments and a careful analysis of transport
distances and possible drift rates of sea ice, transit times of sea ice from some
source areas to the Lomonosov Ridge must have exceeded 1 year, requiring a
perennial ice cover as transport agent. The applied drift rates, however, may
underestimate the potential pronounced variability as recorded from recent
measurements. In 2007, the ice cover drifted at much higher rates in the eastern
Arctic Ocean, and the schooner Tara completed her drift from the northeastern
Laptev Sea to the Fram Strait in approximately 15 months (Gascard et al. 2008).
Thus, even calculations using drift rates of 3 cm sK1 (Darby 2008) may result in
transit times being too long and the CAO might have been easily reached by sea
ice from most potential sediment source areas in less than a year. In this case,
transit times from Canadian Arctic source regions being still longer than a year
may indicate the presence of multi-year sea ice in some regions rather than the
development of a year-round sea-ice cover in the whole CAO. Furthermore,
Phil. Trans. R. Soc. A (2009)
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39
the work by Cronin et al. (2008) disagrees with a perennial ice cover in the CAO
since the Mid-Miocene because abundance maxima of agglutinated foraminifers in
the Early Pleistocene might have been related to seasonally ice-free conditions.
Dowsett (2007) used evidence for a reduced sea-ice cover in the marginal Arctic
Ocean and a warmer terrestrial climate on the circum-Arctic continents to
suggest a seasonally ice-free Arctic Ocean in the Mid-Pliocene (ca 3.3–3 Ma).
Cronin et al. (1993) already proposed that a stronger advection of warmer water to
the Arctic Ocean during warming events recorded in circum-Arctic shallowmarine sediments between 3.5 and 2 Ma might have ultimately led to a reduced
sea-ice cover in the CAO owing to a weakened or eliminated halocline. These
discrepancies between reconstructions may be partly explained by the variable
temporal resolution of the records, leading to a possible under-representation of
rather short interglacials (cf. Funder et al. 2001) in the studied records.
The circum-Arctic shelf seas were seasonally ice covered in the Pliocene, but
the extent of the sea-ice cover is not known. Cronin et al. (1993 and references
therein) concluded from ostracod analysis that the North American continental
margin was ice free in summer during warmer intervals between 3.5 and 2 Ma,
and probably the Arctic Ocean was year-round ice free in some regions at the
Canadian Arctic and Alaskan continental margins. Diatom assemblages from the
marginal seas of the Eurasian Arctic suggest cold conditions and the presence of
a seasonal sea-ice cover since the Late Miocene and probably a sea-ice cover
similar to today since the Late Pliocene (Polyakova 2001). Dinoflagellate cysts in
Mid- to Upper Pliocene sediments from the Yermak Plateau indicate a variable
but still seasonal sea-ice cover (Knies et al. 2002; figure 4). In the Nordic Seas,
sea ice played an important role in the hydrological cycle since the latest Miocene
(Henrich et al. 1989, 2002; Wolf & Thiede 1991). Polar surface waters with a
dense sea-ice cover, at least during wintertime, extended into the Norwegian Sea
in the Late Pliocene (Henrich et al. 2002). The occurrence of sea-ice diatoms
suggests the presence of sea ice in the marginal northwestern North Pacific
Ocean since the Miocene–Pliocene transition (Barron 2003).
The available data suggest strong temporal and spatial variabilities of the seaice cover rather than a stable perennial ice cover in the whole Arctic Ocean. At
present, micropalaeontological data do not prove or disprove the presence of
seasonal or perennial sea-ice cover in the CAO. The ice drift rates, at least in the
Early and Mid-Pliocene, might have been quite different from the average modern
oceanographic situation owing to a more vigorous exchange of water masses with
the North Atlantic Ocean (Henrich et al. 2002; Poore et al. 2006), an increased
input of warm Atlantic waters into the Arctic Ocean that might have additionally
induced a reduction of the sea-ice cover (Butt et al. 2002) and an export of arctic
waters through the Bering Strait until 3.6 Ma (Marincovich 2000). This might
have caused a much faster transport of sea ice from the shelves into the CAO and
an increased export into the North Pacific Ocean and Nordic Seas. Since 3.6 Ma
low-salinity waters might have been increasingly transported from the Pacific
Ocean into the Arctic Ocean, causing enhanced sea-ice formation that increased
albedo and isolated the high heat capacity of the ocean from the atmosphere
(Haug et al. 2001). Evaporative cooling of surface waters during deep-water
formation may have provided the necessary moisture for ice-sheet growth to the
high northern latitudes.
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40
J. Matthiessen et al.
5. Conclusions
Although the Arctic Ocean played a crucial role in the global climate
development in the Pliocene owing to the build-up of ice sheets on the adjacent
continents, relatively few marine data are available to test palaeoceanographic
hypotheses based on numerical and conceptual models. In contrast to other
ocean basins, tectonic forcing and sea-level fluctuations must have strongly
influenced the climate development both on land and in the ocean by changing
the geometry of gateways and by altering the topography of the circum-Arctic
continents. This might have influenced atmospheric circulation and moisture
transport, and subsequently the hydrological cycle of the Arctic Ocean.
A different freshwater supply might have been a decisive factor for triggering
changes in water-mass properties, sea-ice formation and deep convection. These
boundary conditions make the Pliocene Arctic Ocean much different from the
modern basin, but they are rarely considered in palaeoceanographic studies.
Palaeoceanographic reconstructions are generally hampered by low contents of
biogenic components in the high-latitude sediments and sparse palaeo records
from the Arctic Ocean. Despite these restrictions, there is evidence both for a
long-term cooling trend and for pronounced short-term variability. However,
there is still a need for high-resolution studies on time scales that may capture
palaeoenvironmental variability related, for example, to sea-level change or
variable freshwater discharge.
This study was partly funded by the German Research Foundation (STE412/19, MA3913/2-3,
Fi442/10-13) and the Norwegian Research Council. Svend Funder and David McNeil are
acknowledged for their thoughtful reviews.
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