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Geochimicaet CosmochimicaActa, Vol. 61, No. 19, pp. 4181-4200, 1997
Copyright© 1997 ElsevierScienceLtd
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Pergamon
PII S0016-7037(97) 00215-9
Strontium, neodymium, and lead isotope variations of authigenic and silicate sediment
components from the Late Cenozoic Arctic Ocean: Implications for sediment
provenance and the source of trace metals in seawater
BRYCE L. WINTER, CLARK M. JOHNSON, and DAVID L. CLARK
University of Wisconsin, Department of Geology and Geophysics, Madison, Wisconsin 53706, USA
(Received August 15, 1996; accepted in revised form June 6, 1997)
A b s t r a c t - - P r o v e n a n c e changes of silicate sediment deposited during the Late Cenozoic ( 5 - 0 Ma) on
the Alpha Ridge, central Arctic Ocean are determined from variations in strontium, lead, and neodymium
isotope compositions. Whereas strontium and lead isotope compositions are relatively invariant from
~ 5 to 1.7 Ma, end values start to increase at ~ 3 Ma. Subsequently, S7Sr/86Sr and 2°6pb/z°4Pb ratios
progressively increase and end values progressively decrease from ~ 1.7 Ma to the present day. From
these isotope variations, three different endmember compositions for sediment source regions are defined.
The two endmember compositions defined by sediment that was deposited by sea ice from 5 to 1.7 Ma
are consistent with a significant component being derived from the East Siberian Shelf. The progressive
change in isotope compositions from 1.7 to 0 Ma correlates with an increase in coarse detritus deposited
by icebergs over this time period. The isotope data are consistent with a progressive increase in the
proportion of sediment that was deposited in the central Arctic Ocean since 1.7 Ma having been derived
from the northern Canada or Queen Elizabeth Island region.
Changes in the strontium, neodymium, and lead isotope compositions of Arctic seawater are determined
by analyzing the oxide fractions of Fe-Mn micronodules. Strontium isotope compositions of the micronodule oxide fractions are similar to published Late Cenozoic seawater data and indicate that >97% of
the Sr is seawater derived, but minute contributions of Sr from silicate detritus prohibit using the
hydrogenous fractions for chronostratigraphic purposes. Neodymium and lead isotope compositions of
the micronodule oxide fractions, which reflect those of Arctic seawater, follow the isotope variations of
the silicate components throughout the Late Cenozoic. These relations indicate that river water can not
be the primary source of rare earth elements and Pb to the dissolved reservoir in Arctic seawater, and
that ice rafted detritus, by dissolution or exchange processes, is an important source of trace elements
to seawater in ice-covered oceans. Copyright © 1997 Elsevier Science Ltd
1. INTRODUCTION
dymium isotope studies provide important information concerning ocean circulation and the cycling and sources of
trace metals in seawater. Because marine hydrothermal vent
systems act as a net rare earth element (REE) sink (Olivarez
and Owen, 1989; German et al., 1990) and carbonate sediments have very low REE contents (e.g., Palmer, 1985;
Shaw and Wasserburg, 1985; Sholkovitz and Shen, 1995),
both are insignificant sources of Nd to seawater. The source
of the continental input (i.e., via river water or particulate
matter) is the principal parameter that controls the neodymium isotope composition of seawater (Piepgras et al., 1979;
Goldstein and O'Nions, 1981; Goldstein et al., 1984;
Goldstein and Jacobsen, 1987). Recent marine precipitates,
such as foraminifera and Fe-Mn oxides, have neodymium
isotope compositions that are virtually identical to direct
measurements of bottom water masses (Piepgras and Wasserburg, 1980; Palmer and Elderfield, 1985; Staudigel et at.,
1985/86; Albarede and Goldstein, 1992; Jones et al., 1994).
Numerous studies of neodymium isotope variations of preNeogene marine precipitates yield important data for paleogeographic reconstructions and for understanding paleoceanographic circulation (Shaw and Wasserburg, 1985, Staudigel et al., 1985/86; Palmer and Elderfield, 1986; Keto and
Jacobsen, 1987, 1988; Derry and Jacobsen, 1988; Jacobsen
and Plimentel-Klose, 1988; Stille, 1992; Whittaker and
Kyser, 1993). In contrast, few studies have defined the tem-
The isotope variability of Sr, Nd, and Pb in seawater at
a particular geologic time is dependent upon the oceanic
residence time of the element relative to the global ocean
mixing time ( ~ 1 0 0 0 y). Because the residence time of Sr
in seawater is long ( ~ 2 . 5 m.y.; cf. Hodell et al., 1990), the
87Sr/86Sr ratio of seawater is globally uniform (i.e., variability of <1 × 10 -5, Capo and DePaolo, 1992) over timescales of ~ 5 0 k.y., even in relatively enclosed ocean basins
like the Arctic (Winter et al., 1997a). Changes in the strontium isotope composition of seawater over geologic time are
the result of variations in the flux or isotope composition of
Sr delivered to the oceans from continental weathering (i.e.,
river input), mid-ocean ridge circulation (i.e., mantle input),
and dissolution of shelf carbonates (cf. Hodell et al., 1990;
Richter, et al., 1992). Analyses of seawater precipitates,
particularly foraminifera, have defined the strontium isotope
variations of the world ocean throughout the Late Cenozoic
to a high degree of precision (cf. Farrel et al., 1995).
Direct analysis of seawater indicates that the major ocean
basins and even distinct water masses within an ocean have
unique ranges of neodymium isotope compositions (e.g.,
Piepgras and Wasserburg, 1980; Stordal and Wasserburg,
1986; Bertram and Elderfield, 1993); this reflects the short
seawater residence time of Nd ( ~ 1000 years). Marine neo4181
4182
B.L. Winter et al.
poral neodymium isotope variations of a particular ocean
basin during the Late Cenozoic (e.g., Futa et al., 1988; Burton et al., 1994; Stille et al., 1994). Neodymium isotope
variations of modern marine Fe-Mn nodules, which have
slow growth rates and represent accumulation over several
million years (cf. Dymond et al., 1984; Mangini, 1988),
reflect the general patterns of present-day ocean circulation
(Albarede and Goldstein, 1992; Jones et al., 1994). This
observation suggests that despite the rapid ( 1 0 - 1 0 0 k.y.)
and probably dramatic perturbations in ocean circulation that
occurred during glacial periods (Boyle, 1988; Broecker and
Denton, 1989), longer-term water mass movements, and the
continental fluxes and sources of Nd, have remained relatively stable throughout the Pleistocene (Albarede and
Goldstein, 1992).
Because of the low Pb contents in bottom seawater ( - 1
ng/kg), direct measurement of lead isotope compositions is
commonly affected by anthropogenic influences (Schaule
and Patterson, 1981; Flegal and Patterson, t983). However,
marine Fe-Mn oxide nodules, which have high Pb contents
(100-1500 ppm), provide a good measure of the average
lead isotope composition of local bottom seawater over the
growth time of the nodule (Chow and Patterson, 1959, 1962;
Abouchami and Goldstein, 1995). Lead isotope compositions of marine Fe-Mn nodules are variable within an ocean
basin, which is a consequence of the very short residence
time of Pb ( < 100 years; Schaule and Patterson, 1981; Flegal
and Patterson, 1983). As with Nd, recent studies demonstrate that marine lead isotope variations also reflect the
present-day pattern of ocean circulation (Abouchami and
Goldstein, 1995). Lead isotope variations in seawater are
primarily dependent upon the local continental input, but
marine hydrothermal processes can contribute mantle Pb in
the vicinity of relatively fast-spreading ridges (e.g., East
Pacific Rise; Abouchami and Goldstein, 1995).
Strontium, neodymium, and lead isotope compositions of
silicate sediment also provide information about sediment
source regions, including their location, general lithologic
composition, age, and crustal depth of exposure (e.g.,
McLennan et al., 1993; Nakai at al., 1993; Revel et al.,
1996), allowing reconstruction of sediment transport pathways and transport mechanisms (e.g., Grousset et al., 1988;
Gwiazda et al., 1996). These results provide contraints on
paleoceanographic or paleoatmospheric circulation, and ultimately contribute to better paleoclimate reconstructions.
Contrasts in isotope compositions of detrital silicate sediment and those of contemporaneous bottom waters or authigenic sediment components may yield additional information regarding the source of trace metals (i.e., aeolian, hemipelagic, fluvial, or ice-rafted particulates or river water) to
seawater (e.g., Jones et al., 1994; Revel et al., 1996).
In this contribution we present the results of the first strontium, neodymium, and lead isotope study of marine sediments from the Arctic Ocean (Fig. 1 ). From analysis of the
hydrogenous and silicate sediment fractions, we simultaneously track the isotope variations of seawater and detrital
sediment throughout the Late Cenozoic. We place constraints on the timing of changes in sediment source regions
and the geochemical character and general location of the
sediment source areas; this is accomplished by comparing
data from the Arctic Ocean with data from other ocean basins
and with crustal data from circum-Arctic continental masses.
In addition, we address the sources of Nd and Pb in Arctic
seawater.
2. ARCTIC OCEAN
The Arctic Ocean influences global climate through its
effect on the surface heat budget (i.e., ice coverage affects
the local radiation balance as well as ocean-atmosphere heat
exchange) and through export of fresh water, which becomes
incorporated in North Atlantic Deep Water, and thus plays
a role in regulating the intensity of global thermohaline circulation of the world ocean (cf. Aagaard and Carmack, 1994).
The Arctic is a region that is sensitive to global change. In
particular, global climate models indicate that temperature
changes will be enhanced in the Arctic (Sarmiento and Toggweiler, 1984; Broecker and Peng, 1989), and that feedbacks
within the Arctic, which are complexly interconnected between the land, atmosphere, and ocean, may amplify global
climate change. Reconstructing the paleoceanographic and
paleoclimatic evolution of the Arctic Ocean is of critical
importance for understanding the evolution of the modern
global climate and ocean circulation systems (e.g., Clark et
al., 1990).
2.1. Modern Setting
The modern Arctic Ocean is unique in that it is a relatively
enclosed basin and the surrounding continental masses have
crustal ages that range from Early Archean to Cenozoic. The
extreme compositional and age variability of the circumArctic continental crust indicates that changes in the source
region of sediment delivered to the central Arctic Ocean
may produce large changes in the isotope compositions of
sediment and/or seawater over time. Seafloor spreading
along the Nansen-Gakkel Ridge, the northern extension of
the mid-Atlantic Ridge in the Arctic ocean, occurs at an
exceptionally slow rate ( 1 - 3 mm/year; Johnson, 1990),
which implies minimal input of hydrothermal products.
There is no recognized ocean plate subduction at the margins
of the Arctic Ocean, and there is a distinct absence of circumArctic volcanic arc terranes of Tertiary age. These factors
indicate that young, mantle-derived material will not be a
significant component of the sediment deposited in the Arctic
Ocean, in contrast, for example, to the Pacific Ocean.
Surface waters feed the Arctic Ocean through the Bering
Strait ( ~ 5 0 m depth, ~85 km wide) from the north Pacific
(0.8 Sv; 1 Sv = 106 m 3 s -1) and across the Barents Sea
shelf (300-400 m depth) from the northernmost Atlantic
(1.5-2.0 Sv; Carmack, 1990). The modern Arctic Ocean
has deep water communication ( 1 Sv enters and 1.5 Sv exits)
with the lower latitude oceans only through the Fram Strait
( ~ 2 6 0 0 m sill depth, - 6 0 0 km wide; Carmack, 1990). The
Lomonosov Ridge ( ~ 1 5 0 0 m sill depth), which is considered to be a rifted fragment of the Barents shelf, separates
the Arctic Ocean into the Amerasian and Eurasian Basins
( ~ 4 km maximum depth; Johnson, 1990; Fig. 1 ) and greatly
inhibits intraoceanic deep water circulation (Aagaard et al.,
1985; Jones et al., 1995). Deep waters of the Amerasian
Provenance of Cenozoic sediment based on isotope ratios
4183
I Elizabeth
MENDELEYEV
'MARKAROV
RIDGE
BASIN
CA N A
BASIN
'
DA
\
130"
110 °
Fig. 1. Location of Fletcher (FL in Table 2) and Cesar 11 cores in the Arctic Ocean. The dark shaded regions are
continental land masses; the light shaded regions define the continental shelves that are submerged by less than 1000
m of water. The contour interval for the deep ocean basin (unshaded region) is 1000 m. The Lomonosov Ridge
(L.R.) separates the Amerasian and Eurasian Basins of the Arctic Ocean. AR = Alpha Ridge; N-GR = NansenGakkel Ridge; MR = Mackenzie River Delta; B = Banks Island; V = Victoria Island; AH = Axel Heiberg Island;
E = Ellesmere Island; D = Devon Island; Ba = Baffin Island.
Basin are relatively isolated and have ~ 7 0 0 year exchangetimes with the North Atlantic ( O s t l u n d et al., 1987).
River discharge into the Arctic O c e a n from North A m e r i c a
and Eurasia represents ~ 10% of the global river runoff (Aagaard and Carmack, 1989). River water comprises up to
10% of Arctic surface waters ( O s t l u n d and Hut, 1984;
Schlosser et al., 1994). The enclosed nature of the Arctic
Ocean, relatively long deep-water residence times, and the
significant input of river water apparently does not affect the
strontium isotope composition ( W i n t e r et al, 1997a), but
these factors are important influences on the n e o d y m i u m and
lead isotope compositions of Arctic seawater, because of the
short seawater residence time of Nd and Pb.
The central Arctic O c e a n is perennially ice covered, and
present-day ice is predominantly ( ~ 9 9 % ) sea ice that forms
on the wide continental shelves (Clark, 1990). Continental
4184
B.L. Winter et al.
shelves comprise ~35% of the total surface area of the Arctic
Ocean, and they represent ~25% of the total global shelf
area. The large discharge of river water and salt-distillation
effects associated with sea ice formation, result in the development of the cold Arctic halocline ( ~ 5 0 - 2 0 0 m depth),
which produces strong stratification and restricts vertical
mixing in the Arctic Ocean (Aagaard et al., 1981; Rudels et
al., 1996): The cold halocline helps maintain the ice cover
by effectively trapping the oceanic heat of the underlying,
warmer Atlantic layer ( ~ 2 0 0 - 4 0 0 m depth) and deep waters. Perennial sea ice coverage limits the amount of sediment that is deposited by aeolian processes in the deep Arctic
basins (cf. Pfirman et al., 1989, 1990) relative to that deposited in lower latitude, deep ocean basins. Most sediment in
the deep central Arctic Ocean is derived from the shelf regions and is deposited by ice rafting (Clark et al., 1980;
Pfirman et al., 1989; Clark, 1990).
2.2. Stratigraphy and Sedimentology
The lithostratigraphy of the Late Cenozoic ( ~ 5 - 0 Ma)
sedimentary sequence ( ~ 6 m thick) in the central Arctic
Ocean has been reconstructed from several hundred short
piston cores taken along the Alpha-Mendeleyev Ridge
(Clark et al., 1980; Minicucci and Clark, 1983; Mudie and
Blasco, 1985; Morris et al., 1985). This sediment sequence,
which was deposited at very low sedimentation rates ( ~ 1
mm/k.y.), can be divided into three general packages based
on lithology (Fig. 2): (1) a lower preglacial package composed of sediment rafted by sea ice (units A3-A; ~ 5 . 0 - 2 . 4
Ma), (2) a middle transitional package (units A-E; ~ 2 . 4 1.7 Ma), and (3) an upper cyclic package dominated by
sediment that was rafted by glacial ice (units F-M; ~ 1 . 7 0 Ma).
The lower ( ~ 5 - 2 . 4 Ma) sedimentary package is composed of silty mud within which no glacial dropstones have
been found (Clark, 1996); it contains an average of 5 wt%
coarse ( > 6 3 /zm) material, which is dominated by quartz
( 7 0 - 9 8 % ) and authigenic ferromanganese micronodules
( 2 - 30%). The micronodules, which range from ~ 5 0 - 250
/zm in diameter, are aggregates of authigenic Fe-Mn oxides
that cement detrital quartz and clay minerals (Winter et al.,
1997b). The presence of Fe-Mn micronodules and the relatively high degree of bioturbation (Fig. 2) indicate that the
silty mud was deposited at exceptionally low sedimentation
rates ( < 1 mm/k.y.; Clark et al., 1980).
The base of the middle transitional sedimentary package
( ~ 2 . 4 Ma; Fig. 2) is defined by a thin layer of sandy mud
(upper unit A; Clark, 1996) that was deposited by glacial
ice (i.e., icebergs that calve off of continental glaciers and
deposit sediment of all size ranges). This coarse sediment
represents the first major glacial ice surge into the Arctic
Ocean. A second thin layer of sandy mud is present in unit
C (Clark, 1996).
The base of the upper sedimentary package (unit F, ~ 1.7
Ma; Fig. 2) is defined by a significant increase in coarse
( > 6 3 #m) material. The upper sedimentary package ( 1 . 7 0 Ma) is cyclic, composed of alternating sandy and silty mud
intervals. The sandy mud intervals, which have abundant ( 8 25%) dropstones (0.25-3 mm), were deposited by icebergs
|
0 Ma
Biotur%
bation Coarse
m
m
o
"o
0.8 Ma
o
u
9
._w
i
@
o.
w
m
]
m
E
a
0
.u
1.3 Ma
:E
4
t
ij
2.4 Ma
c
o
m
u
0
O.
mm
m
~D
e"o
3 . 4 Ma
o
1:
_8
#J
m
(.O
5.3 Ma
t. n
Silty Mud I
2040
Ill
Fig. 2. Late Cenozoic lithostratigraphy (Clark et al., 1980; Minicucci and Clark, 1983; Mudie and Blasco, 1985; Morris et al., 1985)
and chronostratigraphy (Clark et al., 1986; Jones, 1987) of central
Arctic Ocean sediment (Clark et al., 1980). Ages of the stratigraphic
units are based on extrapolation from magnetic reversals using a
sedimentation rate of 1 mm/k.y. Three sedimentary packages are
defined on the basis of sediment texture: (1) a lower preglacial
package composed of sea ice sediment (units A3-A; ~5.0-2.4 Ma),
(2) a middle transitional package (units A-E; 2.4-1.7 Ma), and (3)
an upper cyclic package dominated by glacial sediment (units F-M;
1.7-0 Ma). Silty mud was deposited at lower sedimentation rates,
as reflected by higher degrees of bioturbation.
during glacial maxima and deglacial periods. Coarse material
increases in abundance from unit F to unit M (Fig. 2; Clark
et al., 1990), which indicates that since ~ 1.7 Ma, a progressively greater proportion of sediment in the central Arctic
Ocean was deposited by glacial ice. Except for the presence
of dropstones in the upper sedimentary package, the silty
mud intervals of the upper sediment package are identical
to the silty mud that composes the lower sedimentary package. Silty mud was deposited primarily during interglacial
periods by sea ice (Clark et al., 1980), which is ~ 2 m
thick ice that forms on the continental shelves and chiefly
transports clay and silt (Pfirman et al., 1990; Reimnitz et
al., 1992).
3. ANALYTICAL TECHNIQUES
Samples for this study were taken from eleven different cores,
which were chiefly obtained from ice-island T-3 as it drifted over
Provenance of Cenozoic sediment based on isotope ratios
4185
Table 1. Location, stratigraphic position, and sample mass data (in mg) for foraminifera, Fe-Mn micronodules and bulk sediments from the
Alpha Ridge, Central Arctic Ocean.
Sample
number
Water
depth
Latitude Longitude (m)
Total
sample
mass
(mg)
Oxide
(calcite)
mass
Silicate
mass
Percent
oxide
(calcite)
Percent
silicate
Stratigraphic
unit
Core (depth)
Age
(Ma)
1-plank
uppermost M
2-plank
uppermost M
2-benthic uppermost M
FL-200 (0-1 cm)
FL-474 (0-1 cm)
FL-474 (0-1 cm)
0.05
0.05
0.05
80 10.55
85 20.88
85 20.88
172 19.63
11000.23
11000.23
3048
1647
1647
27.58
28.30
12.11
22.94
23.96
9.04
4.64
4.34
3.07
83.2
84.7
74.6
16.8
15.3
25.4
middle K
lower K
middle I
middle G
E/F boundary
middle D
lower D
C/D boundary
uppermost A
upper A
A2-A3
middle A1
upper A2
middle A2
lower A2
upper A3
FL-286 (106-109 cm)
FL-286 (112-115 cm)
FL-286 (202-205 cm)
FL-443 (220-221 cm)
Cesarl 1 (207-209 cm)
FL-443 (289-290 cm)
FL-275 (206-207 cm
Cesarll (210-211 cm
Cesarl 1 (225-228 cm
Cesarl 1 (227-228 cm
FL-380 (285-286 cm)
Cesarl 1 (276-277 cm
Cesarll (333-334 cm
Cesarl 1 (395-396 cm
Cesar 11 (446-447 cm
Cesarl 1 (447-448 cm
0.78
0.80
1.00
1.30
1.50
1.70
1.72
1.75
1.80
2.00
?
2.90
3.40
4.10
4.60
4.65
84 00.84
84 00.84
84 00.84
85 57.96
85 50.90
85 57.96
83 30.23
85 50.90
85 50.90
85 50.90
84 37.54
85 50.90
85 50.90
85 50.90
85 50.90
85 50.90
14402.17
14402.17
14402.17
121 07.71
108 21.20
121 07.71
149 58.64
108 21.20
108 21.20
108 21.20
128 27.89
108 21.20
108 21.20
108 21.20
108 21.20
108 21.20
2316
2316
2316
2436
1380
2436
2884
1380
1380
1380
2401
1380
1380
1380
1380
1380
6.10
3.18
6.62
5.09
6.67
6.63
7.04
3.57
1.75
4.11
15.64
3.97
8.78
17.18
54.00
4.73
4.01
2.54
1.47
1.23
2.65
2.46
2.42
2.51
.
2.76
4.26
1.97
3.34
5.65
9.06
1.74
1.75
0.64
5.15
3.86
4.02
4.17
4.62
1.06
.
.
1.35
11.38
2.00
5.44
11.53
44.94
2.99
65.7
79.9
22.2
24.2
39.7
37.1
34.4
70.3
.
67.2
27.2
49.6
38.0
32.9
16.8
36.8
28.7
20.1
77.8
75.8
60.3
62.9
65.6
29.7
uppermost M
uppermost M
uppermost M
middle K
middle K
middle G
lower F
lower D
upper AI
upper A2
middle A2
middle A3
FL-300 (0-1 cm)
FL-508 ( 2 - 3 cm)
FL-530 (0-1 cm)
FL-428 (81-82 cm)
FL-508 (107-108 cm)
FL-428 (205-206 cm)
FL-443 (268-269 cm)
FL-275 (206-207 cm)
FL-284 (254-255 cm)
FL-420 (417-418 cm)
Cesarll (381-382 cm)
Cesarll (497-498 cm)
0.05
0.05
0.05
0.78
0.78
1.30
1.48
1.72
2.64
3.40
3.90
5.10
85 18.53 14402.85
84 07.50 112 13.00
8449.70 099 30.60
86 03.18 134 35.28
84 07.50 112 13.00
86 03.18 134 35.28
85 57.96 121 07.71
83 30.23 149 58.64
83 47.34 145 50.92
8446.59 122 55.14
85 50.90 108 21.20
85 50.90 108 21.20
2082
1866
1985
2271
1866
2271
2436
2884
2681
2248
1380
1380
Foraminifera
Fe-Mn Micronodules
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
32.8
72.8
50.4
62.0
67.1
83.2
63.2
Bulk Sediment
19
20
21
22
23
24
25
26
27
28
29
30
the Alpha Ridge from 1963 to 1973. The sedimentology and paleontology of these cores have been characterized in previous studies at
the University of Wisconsin. After a preleaching sequence and rinsing with ultraclean water, the oxide fractions of handpicked Fe-Mn
micronodules ( 4 - 8 mg; Table 1 ) were dissolved with a mixture of
0.2 M ammonium oxalate and 0.2 M oxalic acid (0.2 M AO-OA) for
2 h in an ultrasonic bath. The nondetrital fractions (calcite + oxide
coating) of handpicked foraminifera were dissolved with 10% acetic
acid (HOAc) after extensive cleaning with 1 M ammonium acetate
and with ultraclean water in an ultrasonic bath. The nondetrital
fractions of two foraminifera samples were analyzed in order to
characterize the recent seawater isotope composition (e.g., Palmer,
1985; Palmer and Elderfield, 1985 ), as Fe-Mn micronodules are not
abundant in units L and M (see Fig. 2) in the Central Arctic Ocean
(Clark et al., 1980). Unleached bulk sediment samples (total in
Table 2) and silicate residue fractions from micronodule and foraminifera separates were dissolved with a mixture of concentrated
HNO3 and HF. Silicate fractions of the bulk sediment samples (silicate in Table 2) were isolated by sequentially leaching bulk sediments with 10% HOAc (4 h) and 0.2 M AO-OA in an ultrasonic
bath before weighing, spiking, and then dissolving with concentrated
HF and HNO3. Further details of the chemical processing can be
found in Johnson and Thompson ( 1991 ) and Winter et al. (1997b).
All isotope ratios were measured on a VG Instruments Sector 54
mass spectrometer in the University of Wisconsin (U.W.) Radiogenic Isotope Laboratory. Lead isotope ratios (n = 50) were measured for each sample in multi-collector, static mode. The average
ratios for NBS-981 (n = 52) were (2o.) 2°6pb/z°Tpb = 1.09453
_+ 14, 2°6pb/2°sPb = 0.46297 +_ 14, and z°6pb/2°4pb = 16.912 _+ 3.
Uncertainties for all standard data are reported as the 2-sigma error
of the mean of all the analyses and represent a good measure of the
external reproducibility of mass spectrometric measurements for this
research. Strontium isotope ratios (n = 120) were measured for each
sample using dynamic multicollector analysis; Rb interference was
continuously monitored at mass 85 and was always negligible.
Twenty analyses of NBS-987 run during this time period yielded
(2o) 875r/86Sr = 0.710232 _ 15. Neodymium isotope ratios were
measured as NdO ÷ for all samples, except the leached bulk silicate
sediment samples (silicate in Table 2), which were analyzed as
Nd+; 120 ratios were measured for each sample using dynamic
multicollector analysis. Cerium, praseodymium, and samarium were
continuously monitored during all NdO + and Nd + analyses and were
not present. Oxygen isotope compositions used for NdO + analyses
are 180/160 = 0.00211 and 1 7 0 / / 1 6 0 = 0.000387. 143Nd/144Nd ratios
(2o') for standards during this time period are: La Jolla NdO + (n
= 9) = 0.511872 _+ 05, U.W. Lab normal Ames NdO ÷ (n = 24)
4186
B . L . W i n t e r et al.
T a b l e 2. Rb-Sr, S m - N d , and lead isotope data for foraminifera, F e - M n m i c r o n o d u l e s , and bulk s e d i m e n t f r o m the A l p h a Ridge, central
Arctic Ocean.
Sample #
Rb
Sr
87Sr/a6Sr
87Rb/86Sr
-117
-70.7
--
-138
----
0.709208
0.718887
0.709204
-0.709201
_+ 08
_+ 10
_+ 08
± 08
-2.45
----
5.52
195
3.20
27.0
92.1
7.10
110
-10.9
135
135
12.9
-138
10.5
98.6
-.
200
127
12.7
-260
8.15
-110
--138
4.86
109
5.26
-108
20.6
78.8
-.
103
-218
26.0
-121
71.8
202
203
-143
143
55.9
-212
-164
-.
267
265
31.8
31.9
298
218
218
248
--232
51.7
216
77.4
77.4
211
233
189
-.
54.0
0.711012
0.724573
0.709941
0.710350
0.719395
0.709716
0.715656
0.715658
0.710936
0.714104
0.714100
0.710326
0.710313
0.715835
0.710179
0.715947
0.709563
.
.
0.715069
0.712640
0.710568
0.710611
0.715389
0.709529
0.709534
0.713077
0.709461
0.709458
0.714730
0.709219
0.715080
0.709009
0.708996
0.714837
0.709021
0.713154
0.709154
.
.
0.716713
_+ 18
_+ 10
_+ 23
_ 14
_+ 10
+_ 16
+ 10
_+ 10
_+ 25
_+ 11
_+ 09
___ 22
± 21
_+ 09
+_ 21
+ 11
± 21
-2.59
0.356
-2.21
0.286
1.58
.
-2.72
2.73
0.668
.
1.89
-1.75
--
_+ 09
± 11
_+ 10
_+ 10
+_ 11
± 09
_+ 09
_+ 10
+ 21
-4- 21
_+ 09
_+ 10
___ 08
_+ 10
_+ 11
_+ 09
_+ 07
_ 34
_+ 10
2.17
1.39
1.16
1.15
2.53
0.108
.
1.29
-.
1.71
0.272
1.46
0.197
0.197
1.48
0.256
1.21
--
± 10
5.52
.
133
129
.
131
135
120
-122
-124
.
112
-142
.
.
111
132
.
127
158
106
106
123
123
145
.
136
140
-.
.
.
0 . 7 2 5 0 6 2 ± 07
3.45
0.725009 _+ 08
2.83
.
.
0.720089 _+ 09
3.00
0 . 7 2 0 3 9 8 _+ 08
2.46
0.723861 ± 09
3.29
0.723849 ± 10
.
0.717787 ± 08
2.86
0.717778 _ 08
-0.717279 ± 08
2.47
.
.
.
0.717408 _+ 09
2.37
0.717385 ___ 09
-0.717057 ± 10
-.
.
.
Sm
Nd
147Sm/la4Nd
143Nd/144Nd eNd(0) 2°6pb/2°npb 2°7pb/2°npb 2°8Pb/2°4pb
Foraminifera
1-plank.
1-residue
2-plank.
2-residue
2-benthic
1.05
4.65
2.99 19.1
1.33
5.72
3.06 16.5
0.508 2.24
0.1365
0.0945
0.1403
0.1123
.
0.512079
0.511993
0.512046
0.512020
.
_+ 06
_+ 08
_+ 12
_+ 14
.
- 11.04
-12.72
-11.69
-12.19
.
---18.655
.
0.1296
0.0967
0.1334
-0.1026
0.1249
0.0989
.
0.1318
0.0958
-0.1285
_+ 15 - 1 2 . 2 1
_+ 09 - 1 4 . 2 8
___ 12 - 1 2 . 4 5
-___ 10 - 1 2 . 6 8
_ 11 - 1 0 . 0 5
_+ 10 - 1 0 . 2 3
.
.
_ 12 - 1 0 . 0 8
_+ 10 - 9 . 4 3
_+ I 0 - 9 . 2 3
_ 10 - 1 0 . 4 5
_+ 10 - 9 . 8 1
_+ 11 - 1 1 . 0 6
+ 15 - 9 . 4 1
+_ 09 - 9 . 3 5
_+ 10 - 9 . 0 3
_+ 11 - 9 . 2 1
_ 09 - 9 . 8 2
± 09 - 9 . 2 5
_+ 14 - 9 . 7 8
.
.
± 13 - 9 . 3 5
+_ 10 - 8 . 8 0
± 11 - 8 . 4 7
_+ 11 - 1 0 . 1 0
~ 09 - 9 . 4 2
.
.
± 08 - 1 0 . 8 0
_+ 09 - 1 0 . 4 9
_+ 07 - 1 0 . 9 5
_+ 08 - 1 0 . 4 8
.
.
_+ 09 - 1 1 . 3 9
_+ 14 - 1 1 . 0 7
_+ 13 - 1 1 . 2 2
__. 09 - 1 0 . 5 5
_+ 08 - 1 0 . 4 0
± 22 - 1 0 . 9 1
18.941
18.794
18.925
18.870
18.722
18.738
18.615
.
18.572
18.649
-18.746
-18.588
18.706
18.627
15.580
--18.503
18.610
.
18.718
18.875
0.1064
0.1322
.
0.1014
0.1331
0.0983
0.1452
.
0.1059
0.1538
0.1064
0.1305
0.1305
0.0963
0.512019
0.511913
0.512007
-0.511994
0.512130
0.512121
.
0.512128
0.512162
0.512172
0.512110
0.512143
0.512079
0.512163
0.512166
0.512182
0.512173
0.512142
0.512171
0.512144
.
0.512166
0.512194
0.512211
0.512127
0.512162
.
0.512092
0.512107
0.512084
0.512108
.
0.512061
0.512078
0.512070
0.512105
0.512112
0.512086
18.859
18.605
.
-18.664
18.293
18.663
.
18.621
18.621
18.581
18.627
---
0.1145
0.1170
0.1060
0.1129
0.1183
0.1042
0.1164
.
0.1172
-0.1066
---0.1045
--
0.512020
0.512029
0.511979
0.511966
0.512064
0.512041
0.512049
.
0.512128
0.512146
0.512134
0.512129
0.512107
0.512087
0.512154
0.512144
_ 09
_+ 09
_+ 09
___ 08
± 09
± 15
+ 08
.
+_ 09
___ 09
___ 08
___ 05
± 14
_ 16
_ 12
± 08
18.625
18.981
19.006
19.024
18.927
18.912
18.936
.
18.777
-18.776
-18.826
-18.891
--
---15.609
---38.673
15.638
15.605
15.669
15.637
15.608
15.617
15.633
38.818
38.914
38.823
38.767
38.826
38.654
38.600
15.588
15.589
-15.636
-15.617
15.605
15.584
15.600
--15.584
15.587
38.520
38.636
-38.692
-38.668
38.638
38.603
38.571
--38.414
38.559
15.570
15.639
38.755
38.661
15.624
15.602
38.646
38.530
-15.595
15.555
15.604
-38.602
38.290
38.675
15.601
15.590
15.578
15.604
---
38.637
38.598
38.702
38.644
---
15.61l
15.655
15.620
15.666
15.658
15.651
15.710
38.538
38.900
38.964
38.983
38.856
38.935
39.083
15.625
-15.583
-15.589
-15.596
--
38.746
-38.739
-38.871
-38.859
--
.
Fe-Mn Micronodules
3-oxide
3-residue
4-oxide
5-oxide
5-residue
6-oxide
6-residue
duplicate
7-oxide
7-residue
duplicate
8-oxide
duplicate
8-residue
9-oxide
9-residue
10-oxide
duplicate
10-residue
ll-ox+res
12-oxide
duplicate
12-residue
13-oxide
duplicate
13-residue
14-oxide
duplicate
14-residue
15-oxide
15-residue
16-oxide
duplicate
16-residue
17-oxide
17-residue
18-oxide
duplicate
18-residue
4.52
6.12
3.45
16.3
3.28
9.00
3.71
.
2.98
3.06
3.06
7.02
.
.
3.80
8.62
4.12
3.45
3.45
5.53
9.07
5.50
.
-8.81
.
.
5.64
5.58
.
3.84
6.01
3.43
5.56
.
4.94
13.1
3.78
8.74
8.74
3.29
21.1
38.2
15.6
77.4
19.3
43.5
22.6
.
13.6
19.3
19.3
33.0
.
23.1
37.5
25.3
15.5
15.5
33.5
45.2
25.4
.
59.3
37.2
.
32.0
25.5
.
22.9
27.3
21.1
23.1
.
28.2
51.5
21.5
40.4
40.4
20.6
5.68
6.82
6.16
6.54
7.33
5.39
6.13
.
6.18
-5.66
29.9
35.2
35.1
35.0
37.4
31.2
31.8
.
31.8
31.8
32.0
32.0
28.0
28.0
34.8
34.8
0.0993
0.1387
0.0983
0.1341
0.1341
0.0995
0.1212
0.1309
.
-0.1429
.
.
.
.
Bulk S e d i m e n t
19-total
20-total
20-silicate
21-total
22-total
22-silicate
23-total
duplicate
24-total
duplicate
24-silicate
duplicate
25-silicate
duplicate
26-silicate
duplicate
4.82
-6.03
-12.19
-12.02
-12.99
-13.26
-11.34
-11.78
-11.62
.
-10.10
-9.74
-9.97
-10.07
-10.50
-10.89
-9.58
-9.78
.
Provenance of Cenozoic sediment based on isotope ratios
4187
Table 2. (Continued)
Sample #
27-total
duplicate
27-silicate
duplicate
28-silicate
duplicate
29-silicate
duplicate
duplicate
3G-silicate
duplicate
Rb
Sr
123
.
116
-138
-108
.
.
118
.
137
.
157
153
129
129
189
.
.
153
.
875r/S6Sr
0.714057
.
.
0.713772
0.713784
0.722878
0.722878
0.715659
.
.
.
.
0.716293
.
.
___09
.
± 07
± 08
± 09
± 09
± 09
.
.
± 08
.
87Rb/s6Sr Sm
2.59
6.72
2.14
.
3.10
-1.65
5.28
.
6.15
-4.48
2.24
4.65
Nd
34.2
34.2
30.2
.
36.0
36.0
26.5
26.5
26.5
27.4
27.4
147Smf'~Nd 143Nd/144Nd eNd(0) 2°6pb/2°4pb 2°Tpb~°4Pb 2°8pbfl°4pb
0.1186
0.512232 _ 08 -8.07
-0.512234 + 09 -8.03
0.1057
0.512187 ± 08 -8.94
.
.
.
.
.
0.1030
0.512055 ± 11 -11.52
-0.512058 ± 07 -11.46
0.1022
0.512023 ± 08 -12.14
-0.512025 ± 08 -12.10
-0.512012 ± 07 -12.36
0.1023
0.512097 ± 10 -10.70
-0.512101 ± 09 -10.63
18.556
-18.871
15.633
-15.635
38.531
-39.009
19.001
-18.560
--18.647
--
15.618
-15.575
--15.563
--
39.016
-38.835
--38.708
--
In-run uncertainties given for strontium and neodymium isotope ratios are 2-sigma errors. In-run uncertainties (2-sigma errors) for 2°6Pb/
2°4pb, 2°7pb/zC~pb, and z°spb/2°4pb ratios are less than 0.012, 0.015, and 0.048, respectively. Strontium, rubidium, samarium, and neodymium
concentrations are in ppm. Strontium and neodymium isotope ratios were corrected assuming linear mass fractionation and s651"/88Sr = 0.1194
and 46Ndf44Nd = 0.7219. Lead isotope ratios were corrected for mass fractionation of +0.12% per a.m.u.
eNd(0) = [(143Nd/144Ndsampl~)/(Vt3Nd/144Ndch.uR)- 1]*1000; 143Nd/l't4NdcHua = 0.512645.
plank = nondetrital fraction of planktonic foraminifera separates; benthic = nondetrital fraction of benthic foraminifera separates; oxide
= oxide fraction of Fe-Mn micronodule separates; residue = silicate residue fraction of foraminifera and micronodule separates; 11-ox+res
is the total dissolution (oxide + silicate fractions) of a micronodule separate; total = bulk, nonleached sediment; silicate = silicate fraction
of bulk sediment samples. Duplicates are repeated mass spectrometric analyses of a different aliquot of the same sample solution. Dashes
indicate samples that were not analyzed. See Table 1 for location and stratigraphic information for each sample.
= 0.512150 ± 03, U.W. Lab normal Ames Nd ÷ (n = 16)
= 0.512147 ± 03, BCR-1 NdO ÷ (n = 5) = 0.512645 ± 05. All
procedural blanks were negligible (Pb < 250 pg, Sr < 120 pg, and
Nd < 160 pg). Additional analytical details are given in Table 2.
4. TEMPORAL ISOTOPE VARIATIONS,
CENTRAL ARCTIC OCEAN
Strontium isotope compositions of Fe-Mn micronodule oxide fractions from the Arctic Ocean (0.7090-0.7110; average
= 0.70986) are considerably less radiogenic and markedly less
variable than the silicate fractions (0.7131-0.7251; Fig. 3).
The least radiogenic 87Sr/86Sr ratios of micronodule oxide fractions are consistent with those of Arctic foraminifera and the
reported strontium isotope compositions of lower latitude seawater during the Late Cenozoic (0.7089-0.7092; Farrell et
al., 1995; Winter et al., 1997b); this demonstrates the success
of our chemical leaching technique in concentrating the hydrogenous portion of the micronodules for neodymium and
lead isotope measurements. Simple mixing calculations indicate that greater than 97% of the Sr in the oxide fractions is
seawater-derived. However, minute contributions ( 2 - 3 % ) of
very radiogenic Sr from silicate material, introduced during
diagenesis or our laboratory leaching procedure, results in a
significant increase (~0.0002) in the 87Sr/86Sr ratio of the
micronodule oxide fractions (Winter et al., 1997b). Strontium
isotope variations of the micronodule oxide fractions, therefore, cannot be used to further refine Arctic chronostratigraphy
(Winter et al., 1997b).
Except for one sample (#28, Table 2), 87Sr/a6Sr ratios of
the silicate fractions are relatively uniform ( ~ 0 . 7 1 5 ) from
- 5 . t ) to 1.5 Ma (Fig. 3a). The progressive and dramatic
increase in 875r/86Sr ratios of the silicate fractions from ~ 1.7
Ma to the present-day (0.715-0.725; Fig. 3a) indicates that
a different source region supplied a progressively increasing
proportion of the sediment to the central Arctic Ocean over
this time period.
Temporal lead isotope variations of the hydrogenous and
silicate components of Arctic sediment (Fig. 3b) are similar
to the strontium isotope variations of the silicate fractions,
in that 2°6pb/E°4pb ratios are relatively invariant from ~ 5 1.7 Ma ( 18.63 ), followed by a progressive increase ( 18.6319.03) during the time period of glacial sedimentation, beginning at ~ 1.7 Ma. 2°6pb/2°4pb ratios of the different silicate
fractions are more scattered, but they have the same general
trend as the oxide fractions (Fig. 3b).
Silicate fractions of Arctic micronodules commonly have
eNd values that are slightly lower ( ~ 0 . 2 - 0 . 8 epsilon units)
than those of the corresponding oxide fractions, although
nearly all sediment components, including the bulk silicate
samples, have essentially the same end value (i.e., < __ 0.8
epsilon units) at a given stratigraphic horizon throughout the
Late Cenozoic (Fig. 3c). Two samples of the nondetrital
foraminifera fractions, which are more radiogenic than the
silicate fractions by ~ 1.3 epsilon units, are a possible exception (Fig. 3c). The eNd value of the silicate fraction of one
micronodule sample (3, Table 2) is much lower than samples
of similar age (Fig. 3c) and will not be considered in the
discussions below.
The temporal changes in end values since ~ 1 . 7 Ma (Fig.
3c) are remarkably consistent with the temporal changes in
strontium and lead isotope ratios over this time period and
provide strong support for coherent temporal shifts in source
regions supplying sediment to the central Arctic Ocean. Because lead and strontium isotope ratios are typically inversely correlated with end values in genetically-related
crustal rocks (cf. Faure, 1986), a decrease in end values
accompanying increasing 87Sr/S6Sr and 2°6pb/2°4pb ratios of
Arctic sediment sources is expected. However, in contrast
to strontium and lead isotope compositions, end values of
Arctic siliciclastic fractions are not invariant from - 5 - 1 . 7
Ma, but rather end values increase significantly beginning at
~ 3 Ma (Fig. 3c).
4188
B.L. Winter et al.
A.
r
'i
1
•
3
B.
0
0
.
.
.
l
.
.
.
.
.
.
.
.
q
t
5
X '
'
/
'
0 00~o ~
"Glacial
2
.
(
4
. . . .
.
.
.
-
0
olo °nG'c'at
1
.
.
.
.
,
6
0.705 0.710 0.715 0.720 0.725 0.730
87Sr/86Sr
-
-
-
Q
+
~Non-Glacia~
t-
5"
6
!
18.2
i
|
i
18.4
i
18.6
I
I
18.8
19.0
I
19.2
206pb1204pb
C.
0
l Symbol Key
1
I
Fe-Mn Micronodules
• OxideFraction
O Silicate Fraction
• Total Micronodules
2
Foraminifera
• Non-DetritalFraction
[ ] DetritalFraction
4
Bulk Samalm
X BulkSediment (No.-L~ach.c
"l- BulkSilicate Fraction
5
6
-16
-14
-12
-10
-8
-8
E l ~ i l o n Nd
Fig. 3. Isotope compositions (Sr, Pb, and Nd) of sediment components from the central Arctic Ocean throughout
the Late Cenozoic. The analytical errors are smaller than the symbols. The thick grey curves represent the average
isotope compositionsof the Arctic Ocean sediment during the past 5 m.y. The horizontaldashed lines are the boundaries
of the three sedimentary packages depicted in Fig. 2. The nondetrital fraction of foraminifera is calcite + Fe-Mn
oxide coatings.
5. LOCATIONS OF ARCTIC SEDIMENT
SOURCE REGIONS
Three isotopically distinct source regions for sediment to
the central Arctic Ocean during the Late Cenozoic are inferred from the strontium, neodymium, and lead isotope
compositions (Fig. 3). The approximate composition of the
endmember sources, which probably themselves represent
mixtures of several crustal sources, are depicted in the correlation diagrams of Fig. 4. The source for the oldest sediments
(Source 1, - 5 - 3 . 5 Ma; Fig. 4) is estimated to have end
- 1 3 , 87Sr/a6Sr ~ 0.714, and z°6pb/2°4pb ~ 18.5. We interpret a second source (Source 2; Fig. 4) to be dominant at
~ 3 Ma that has similar strontium and lead isotope compositions as Source 1, but a higher end value of approximately
- 8 . A third source region (Source 3; Fig. 4) is interpreted
to become progressively more important in supplying sediment to the central Arctic Ocean from ~ 1.7 Ma to the present-day. This third source has the most radiogenic strontium
and lead isotope ratios (0.725 and 19.1, respectively) and
comparatively low end values of approximately - 1 3 . The
isotope compositions of Source 3 indicate that the corresponding crustal source rocks have a more supracrustal composition (i.e., high U/Pb and Rb/Sr and low Sm/Nd) and
an older average age relative to source terranes 1 and 2.
5.1. Glacial Marine Sediment ( - 1 . 7 - 0 Ma)
We interpret the progressive change in isotope compositions of the silicate sediment fractions since ~ 1.7 Ma (Fig.
3) to be directly linked to the progressive increase in abun-
Provenance of Cenozoic sediment based on isotope ratios
A,
4189
B.
19.1 L
-8"
o;;o"
,Q
Z
r,
o
m~
-lO -I"/°
==
U,I
+
-14
0.708
'
'
'
'
'
0.712
'
0.716
'
'
0.720
'
'
"
'
i:::rl
'
0.724
0.728
0.708
•
,
,
,
,
0.712
.
,
,
,
,
0.716
,
.
0.720
.
.
,
,
0.724
.
,
0,728
87Sr/8%r
87Sr/86Sr
-6, C .
I
i
=
l
r
l
l
i
l
8F
•
-IOI-
Fe-Mn Ml=ronodulos
Oxide Fraction
0 Silicate Fraction
• Total MIcranoduios
Fommlmlfmm
O=
"-- +
•
Non-Oeh'itel
Fraction
[ ] Debitai Fraction
Bulk flamnlm=
"12 f
-14
18.4
X Bulk Sediment ~==~=)
-~- Bulk Silicate Fraction
I
18.5
I
I
18.6
I
I
18.7
I
I
18.8
I
i
18.9
,
i
19.0
,
i
19.1
206pb/204pb
Fig. 4. Diagrams showing the correlations of lead, neodymium, and strontium isotope compositions for Late
Cenozoic sediment components from the central Arctic Ocean. The thick dashed arrows illustrate the general change
in isotope compositions with depositional age of the sediment and depict three main endmember compositions. Source
#1 reflects the endmember composition of the oldest (5 Ma) sea ice sediment. Source #2 reflects the endmember
composition of sea ice sediment at ~ 3 Ma. Source #3 is the endmember composition of the glacial sediment ( 1.50 Ma) in the Arctic Ocean. The nondetrital fraction of foraminifera is calcite + Fe-Mn oxide coatings.
dance of coarse glacial ice-rafted detritus (IRD) deposited
during this time period (Clark et al., 1990; Fig. 2). During
Pleistocene glacial maxima, North American ice sheets extended to the Arctic coastline from northwest Canada (i.e.,
the northern Yukon District) to Greenland, covering most
of the Queen Elizabeth Islands (QEI), but there is no evidence for glaciation north of the Brooks Range in coastal
Alaska (Fig. 5; Dyke and Prest, 1987a,b; Hodgson, 1989).
The Eurasian ice sheet extended from northern Europe to
the Talmyr Peninsula during the Pleistocene (Velitchko et
al., 1989; Fig. 5), but was probably not an important source
of IRD to the Amerasian Basin, inasmuch as calved icebergs
would have been exported directly through the Fram Strait
by the Transpolar Drift (Bischof et al., 1996; Fig. 5). Eastern
Siberia is thought to have been largely ice-free during the
Pleistocene (Velitchko et al., 1989).
In contrast to the present-day dominance of sea ice, icebergs derived from the northern Canada and QEI regions
composed the major component of Arctic Ocean ice coverage during Pleistocene glacial maxima (Clark, 1990; Bischof
et al., 1996). Continental ice-sheet reconstructions indicate
that ice-divides apparently extended through central Ellesmere Island to northwestern Devon Island, and the M'Clintock Ice Divide extended south from northeastern Victoria
Island to approximately latitude 60°N (Dyke and Prest,
1987a; Fig. 5). Source regions of Pleistocene glacial ice
rafted debris (IRD) to the Arctic Ocean are, therefore, restricted to the north and west of these divides. The Northern
Canada/QEI region was verified as a source of glacial IRD
on the basis of a higher frequency of sandy mud intervals
in Pleistocene sediment cores from the eastern Alpha Ridge
relative to cores from the westem portion of the ridge (Minicucci and Clark, 1983).
Recent detailed analysis of the compositional variations of
coarse ( > 2 5 0 #m) Pleistocene IRD from the central Arctic
Ocean and of glacial till from Arctic North America indicate
that the primary source region for glacially ice rafted quartz
was Axel Heiberg and Ellef Ringes Islands (i.e., the
Sverdrup Basin), whereas glacially ice rafted carbonate was
primarily derived from Banks and Victoria Islands (Darby
and Bischof, 1996; Bischof et al., 1996). Chemical compositions of detrital iron oxide grains from central Arctic Ocean
sediment and from Arctic North American tills also identifies
these two regions, which were covered by the Laurentide
(Banks-Victoria Islands) and Innuition (Axel Heiberg-Ellef
Ringes Islands) ice sheets, respectively, as important source
areas (Darby and Bischof, 1996).
Although no trace element or isotope geochemical data
4190
B.L. Winter et al.
Fig. 5. Long-term, general surface currents of the modem Arctic Ocean as determined from sea ice drift paths
(Gordienko and Lationov, 1969). The anticyclonic Beaufort Gyre (BG) over the Canada Basin and the Transpolar
Drift (TPD) are the two main fields that characterize Arctic surface currents. The East Greenland Current (EGC)
and the Norwegian-Atlantic Current (NAC) are the main systems south of the Fram Strait. The thick hashed lines
define the extent of continental ice sheets during the last glacial maximum (Dyke and Prest, 1987a,b; Velitchko et
al., 1989). LIS = Laurentide Ice Sheet, IIS = Innuition Ice Sheet, GIS = Greenland Ice Sheet, and EIS = Eurasian
Ice Sheet. The dark lines within the ice sheets are the major ice divides and define the general ice flow directions
(Dyke and Prest, 1987a).
exist for marine sediments from the northern Canada/QEI
source region, we suggest that the endmember composition
defined by the youngest Arctic Ocean sediments (i.e., Source
3, Fig. 4) is representative of the QEI region. The progressive
change in the isotope compositions of central Arctic Ocean
sediments since ~1.7 Ma (Fig. 3), and the inferred age and
lithology of the QEI region (see below) are consistent with
this interpretation.
5.2. Sea Ice Sediment ( ~ 5 - 1 . 5 Ma)
Most of the sea ice in the modem Arctic Ocean is formed
on the surrounding continental shelves (Colony and Thorndyke, 1985), where silt and clay are entrained primarily by
suspension-freezing processes (Reimnitz et al., 1992), and
we infer that similar processes played a role in the origin of
the sediment deposited in the deep Arctic Ocean from ~ 5 1.7 Ma (Clark, 1990; Clark et al., 1990). Primarily because
of the lack of coarse material, there has been only limited
progress in determining the provenance of sea ice sediment
in the deep Arctic Ocean. The general pathways of modem
sea ice drift are well-known and consist of the anticyclonic
Beaufort Gyre in the Canada Basin region, and the various
branches of the Transpolar Drift across the Eurasian Basin
(Gordienko and Laktionov, 1969; Fig. 5). The details and
variability of Arctic surface currents, however, are poorly
known. Many global climate models suggest that the large
continental ice sheets of Pleistocene glacial maxima significantly altered atmospheric circulation (Manabe and Broccoli, 1985; COHMAP, 1988), and may have produced a
cyclonic gyre in the Arctic Ocean, which would have reversed sea ice circulation relative to the present day (cf.
Pfirman et al., 1989). Prior to continental glaciation, when
the lower sedimentary package was deposited ( ~ 5 - 2 . 4 Ma;
Fig. 2), it is possible that surface currents were broadly
similar to that of the modem Arctic Ocean.
The Laptev Sea region (Fig. 1) is known to have the
largest modem annual sea ice export rate (500,000 kin2;
Nurnberg et al., 1994) and clay mineralogic data for sediment entrained in sea ice within the Transpolar Drift over
the Eurasian Basin suggest that the Laptev Shelf is the major
source of sediment in modem sea ice over this part of the
Arctic Ocean (Nurnberg et al., 1994). However, insufficient
data currently exist to quantitatively estimate the percent
contribution of sediment from distinct shelf regions that is
entrained within sea ice at various locations in the Arctic
Provenance of Cenozoic sediment based on isotope ratios
Ocean, particularly in the Canada Basin region. The narrow
shelf and the compressional regime created by the northeasterly prevailing winter winds greatly restricts the amount of
sediment transported by sea ice from the Beaufort Shelf
region to the central Arctic Ocean (i.e., in comparison to
the dilational regime in the Laptev Sea; Reimnitz et al.,
1994). Nonetheless, significant amounts of sediment, in
temls of the Beaufort Shelf sediment budget, are probably
transported to Arctic deep marine environments via sea ice
(Kempema et al., 1988). The Chukchi and East Siberian
Seas export minor amounts of modem sea ice in comparison
to the Laptev Sea (cf. Nurnberg et al., 1994), but these
very wide continental shelves are probably also significant
sources of sea ice sediment to the Canada Basin. Although
no isotope geochemical data exist for sediment from these
shelf regions, we hypothesize that endmember sediment
Sources 1 and 2 defined by the older ( ~ 5 - 2 . 4 Ma) sediment
from the Alpha Ridge (Fig. 4) are representative of sediment
from either the Beaufort, Chukchi, or East Siberian Shelf
regi3ns. We base this hypothesis on the general sea ice
movement pathways in the modem Arctic Ocean (Fig. 5).
6. ISOTOPE CHARACTERISTICS OF CRUSTAL
RESERVOIRS AND RECENT GLOBAL SEDIMENTS
The isotope compositions of Arctic Ocean sediment are
pertinent to global geochemical models that have been previously published based on data obtained exclusively from
lower latitude oceans. Sr-Nd isotope compositions of Arctic
Ocean sediment plot along a mixing array that lies between
endmembers defined by young oceanic volcanic rocks and
largely Proterozoic-age upper crest (Fig. 6a); the two low
strontium isotope endmember compositions (Sources 1 and
2; Fig. 4) cannot be resolved at the scale of Fig. 6a. Isotope
compositions of Pacific marine sediments also define a general mixing array between young volcanic rocks (circumPacific arcs) and Proterozoic upper crustal material (Fig.
6a); the latter of which is primarily aeolian dust derived
from the Central Loess Plateau of China (Nakai et al., 1993;
Jones et al., 1994). Marine silicate sediments from the Atlantic Ocean have isotope compositions that plot along mixing
arrays between two and three endmember compositions (Fig.
6a), which include young volcanic rocks (oceanic islands),
Proterozoic upper crest that is probably derived from Caledonian terranes of the British Isles (i.e., represented by sedimeres from the Bay of Biscay), and Archean upper crest
derived from the Superior Province and Baffin Bay regions
(Revel et al., 1996).
Strontium and neodymium isotope data for suspended sediment load from numerous rivers that drain diverse areas
were interpreted by Goldstein and Jacobsen (1988) to define
a mixing relation between young oceanic volcanic terranes,
which have a strong mantle component, and old Archeandominated upper crust. The mixing curve defined by the
river data (Goldstein and Jacobsen, 1988) was proposed
to be representative of the global crest that is exposed to
weathering, and from this relation an average composition
of the upper crest was calculated at end = --16.7, TNd(DM)
= 2.1 Ga, and 875r/86Sr = 0.716.
The average neodymium isotope composition and model
4191
age of Late Cenozoic sediment from the Arctic Ocean (eNd
= --10.7; TNd(DM) = 1.5 Ga), which are similar to those of
loess deposits, river loads, and other young marine sediments
(Fig. 6a), probably reflect the preferential erosion of
younger orogenic belts relative to old stable cratons (e.g.,
Goldstein et al., 1984; Goldstein and Jacobsen, 1988). In
the specific case of Arctic marine sediments, the fact that
end values are higher than those of average upper crust suggests that Archean terranes are unlikely to be major direct
sources to the central Arctic Ocean (Fig. 6a,c). Modem
marine sediment that is dominated by Archean sources is
only apparent in the North Atlantic Ocean; where sediment
has been traced to the Superior Province, western Greenland,
and Baffin Bay (Revel et al., 1996).
Lead isotope compositions of young oceanic volcanic
rocks and Precambrian upper crustal terranes overlap and
are highly variable (Fig. 6b,c), producing Pb-Sr and Nd-Pb
isotope variations that are less distinctive than Nd-Sr isotope
variations (Fig. 6a). We, therefore, suggest that it is misleading to interpret apparently linear Pb-Nd isotope correlations for suspended sediment from selected rivers as representative of the global crest (Asmerom and Jacobsen, 1993)
and the Arctic sediment data presented here demonstrates
that there is no single global Nd-Pb mixing array. We note
that not all river samples analyzed by Asmerom and Jacobsen (1993) are consistent with a linear correlation, and relatively few of the river samples compared to the Goldstein and
Jacobsen (1988) database were measured for lead isotope
compositions. Sample bias for the Pb-Nd isotope data of
river suspended sediment seems likely, given the fact that
2°6pb/2°4pb ratios of sediment from the Philippines rivers
(Pampagna and Agno Rivers), which Asmerom and Jacobsen ( 1993 ) suggest are samples representive of young volcanic rocks (i.e., the mantle), are very unusual in comparison
to the average lead isotope composition of young volcanic
arc terranes (average 2°6pb/2°4pb ~ 18.6; Fig. 6b,c). It seems
likely that silicate sediment from restricted regions of ocean
basins will define distinct Pb-Nd and Pb~Sr arrays (e.g.,
West Central Atlantic data: Fig. 6c; Arctic data: Fig. 6b).
However, the lead isotope variability of Precambrian upper
crustal and young oceanic volcanic reservoirs (Fig. 6b,c),
makes it highly unlikely that the upper crust as a whole
can be described by a general mixing curve involving lead
isotopes.
6.1. Implications for Source Regions to the Arctic
Ocean
On the basis of relations in Fig. 6a, we hypothesize that
the isotope compositions of sediment from the Bay of Biscay
(Fig. 6a) are broadly similar to those of sediment derived
from the northern Canada/QEI region (i.e. Source 3; Fig.
4). Arctic sediments derived from Sources 1 and 2 (Fig. 4)
have neodymium and strontium isotope compositions that
are very similar to those of the deposits from the Central
Loess Plateau of China (Fig. 6a). Although the origin of
these loess deposits is still unknown (cf. Liu et al., 1994),
it is possible that they may approximate the average isotope
composition of the upper crust that was exposed to weathering in Asia during the Late Cenozoic. Lead isotope composi-
4192
B.L. Winter et al.
A.
10
,
,
,
,
,
,
19.8
,
0.3
0.5
0
~
B.
-10
1.0 Z
19.3
"S •[ i
18.8
2.0 ¢~
1•-20
.
.
.
.
2.5 ~
9O08 ~
.
18.3
-30
3.0
-3.5
-40
0.700
I
I
0.710
I
I
I
I
0.720
I
0.730
17.8
0.700
1.740
0.710
87Sr/86Sr
10
0.720
0.730
87Sr/86Sr
C.
Z
&
g
-10
o
~l. -20
I.g
-40
17.8
18.3
18.8
19.3
19.8
2O6pb/204pb
Fig. 6. Isotope compositions of silicate sediments from the Arctic Ocean (open circles). Additional data include:
lower latitude marine silicate sediments (data fields), loess deposits (data fields), and samples of suspended sediment
load from rivers that drain a variety of tectonic terranes that have diverse ages (open squares; Goldstein and Jacobsen,
1988; Asmerom and Jacobsen, 1993). The correlation line for the river data is also given. A.U.C. = Average Upper
Crust as estimated from river data (Goldstein and Jacobsen, 1988; Asmerom and Jacobsen, 1993). Neodymium
depleted-mantle model ages, which were calculated by assuming average crustal 147Sm/144Ndratios (0.11 ), represent
the average mantle extraction age of crustal components in the sediment source regions. The Oceanic Volcanics field
represents undifferentiated volcanic arc and mid-ocean ridge compositions. The sources for the data are as follows:
North Atlantic - Revel et al., 1996; Pacific - Nakai et al., 1993; West Central Atlantic - White and Dupre, 1985;
Atlantic, Pacific, and Indian data - Ben Othman et al., 1989; Loess data - Taylor et al., 1983; Goldstein et al., 1984;
Liu et al., 1994; ; Baffin Bay data - Revel et al., 1996. The Superior Province field is based on composite rock and
sediment samples from McCulloch and Wasserburg (1978) and suspended sediment from the Whale River (Goldstein
and Jacobsen, 1988). Sediment samples from Biscay Bay (cf. Revel et al., 1996) depict the composition of Caledonian
crust in this region. Data fields for Oceanic Volcanics are from: Hegner and Tatsumoto, 1987; White et al., 1987;
Bullen and Clynne, 1990; Leeman et al., 1990; Stern et al., 1993.
tions of the older sea ice sediments from the Arctic Ocean
(i.e., from Sources #1 and #2; Fig. 4) are significantly less
radiogenic than those of sediments from the Indian or Atlantic Oceans (Fig. 6c) and are more similar to the lead isotope
compositions of Pacific sediments. The relatively nonradiogenic lead isotope compositions of sediment from the Pacific
probably reflect a strong mantle component from surrounding young volcanic arcs a n d / o r metalliferous sediment
derived from hydrotherrnal processes associated with the
fast-spreading East Pacific Rise. The Arctic sediments are
not metalliferous, and it is very unlikely that hydrothermal
processes associated with the Nansen-Gakkel Ridge in the
Arctic Ocean could explain the nonradiogenic lead isotope
ratios, because this ridge is spreading at an extremely slow
rate ( ~ 2 rnm/year), and it is isolated from the Amerasian
Basin by the Lomonosov Ridge. W e interpret the relatively
low 2°6pb/2°4pb and 2°Tpb/2°4Pb ratios of endmember Sources
#1 and #2 (Fig. 6b,c; see Fig. 8) to indicate that lower crust,
which is commonly depleted in incompatible elements as a
result of high-grade metamorphism, is an important component in these crustal source regions, as opposed to a young
volcanic arc component.
7. ISOTOPE COMPOSITIONS OF ARCTIC-AMERASIA
CONTINENTAL TERRANES
The continental terranes of northern Siberia, north Alaska,
and the Queen Elizabeth Island ( Q E I ) region are very corn-
Provenance of Cenozoic sediment based on isotope ratios
15.8
O.
15.7
o
a.
o
r~
t5.6
15.5
18.0
18.5
lg.O
19.5
206pb1204p b
Fig. 7. Lead isotope compositions of silicate fractions of Arctic
Ocean sediment. Symbols are as in Figs. 3 and 4. Also shown is the
field for the hydrogenous component (oxide fraction) of Arctic FeMn micronodules (see Fig. 8). Average compositions for Phanerozoic orogenic belts are from Table 3, Caledonian orogenic terranes
that were built on the Precambrian Baltic Shield have very radiogenic
lead isotope ratios that plot off of this diagram. Caledonian orogenic
terranes that were built on or adjacent to crust that was depleted in
U during high-grade metamorphism have less radiogenic lead isotope
ratios. Average lead isotope compositions for the Caledonian terranes of Greenland, Scotland, Ireland, and Norway (GSIN) are very
similar. The field for North Atlantic sediments is from Ben Othman
et al. (1989). Stacey and Kramers ( 1975 ) model for the lead isotope
evolution of average crust is shown for reference (S-K).
plex and diverse (cf. Trettin, 1989; Zonenshain et al., 1990;
Moore et al., 1994). Each of these regions consists of extensive Paleozoic and Mesozoic composite terranes that developed on, and adjacent to, Archean and Proterozoic cratons.
The composite terranes consist of orogenic fold belts, rift
complexes, intrusive complexes, and thick clastic sequences
derived from Archean shields and Proterozoic to Mesozoic
orogenic highlands (cf. Trettin, 1989; Zonenshain et al.,
1990; Moore et al., 1994). In an attempt to delineate crustal
regions as potential sources of sediment to the Amerasian
Basin during the Late Cenozoic, we have compiled published
neodymium and lead isotope data for whole-rock samples
from a variety of continental terranes (Table 3; Fig. 7).
Isotope data for modern riverine and adjacent shelf sediments would best characterize the average composition of
these crustal terranes, but unfortunately, data for such samples from the Arctic do not yet exist. We have focused our
Arctic crustal compilation (Table 3) on whole-rock lead
isotope data, particularly on Phanerozoic hydrothermal ore
deposits, which scavenge Pb from large regions of the crust
and, therefore, decrease the effects of local heterogeneities.
Comparisons of our isotope data on Arctic marine sediments
with crustal isotope data is hampered by the absence of
published data from northeast Siberia and the QEI region.
The major features of the northeast Siberian crust include
the high-grade Archean and Proterozoic basement rocks and
overlying Proterozoic platform sediments of the Siberian
Craton, as well as Paleozoic and Mesozoic composite terranes of the Verkhoyansk, Kolymian, Anyui, and Chukotka
fold belts that lie east of the Siberian Craton (e.g., Fujita
and Newberry, 1982; Zonenshain et al., 1990; Rundqvist
4193
and Mitrofanov, 1993). Much of the exposed crystalline
basement of the Siberian Craton consists of granulite-facies
metamorphic rocks that are similar to those of the wellknown Early Archean terranes of southwest and southeast
Greenland (Rundqvist and Mitrofanov, 1993 ). The existence
of widespread high-grade basement in the Siberian Craton
is supported by low 2°rpb/2°4pb ratios of crustally-contaminated basalt from the Siberian Traps (e.g., Sharma et al.,
1992; Lightfoot et al., 1993; Wooden et al., 1993; Table 3);
although, it is unclear if such a component is widely exposed
at the surface. The only lead isotope data (n = 5; Table
3) for northeast Siberia that have been reported in western
literature are from ore deposits within the Late Paleozoic to
Jurassic Verkhoyansk Fold Belt (Karpenko et al., 1981).
This fold belt, which has clastic components that were primarily derived from the Siberian Craton, is the most extensive composite terrane of northeast Siberia, and hence, may
be an important source of sediment to the Amerasian Basin.
Lead isotope compositions of these ore deposits are quite
variable (Table 3; Karpenko et al., 1981), but four of the
five data are relatively nonradiogenic (z°rpb/2°4pb = 18.118.8) and may indicate a mixture of high-grade crust (i.e.,
Siberian basement) and Phanerozoic orogenic sedimentary
rocks. These data are consistent with the nonradiogenic Pb
component of Sources 1 and 2 (Fig. 4) for the older ( ~ 5 1.7 Ma) Arctic sea ice sediments that we interpret to reflect
input from the Siberian Craton.
The Chukotka (New Siberian-Chukchi) Foldbelt is the
northernmost composite terrane of northeast Siberia and is
•also a probable source of sediment to the adjacent Arctic
shelf region and ultimately to the deep Amerasian Basin.
The Chukotka Foldbelt is considered by many workers to
be part of an Arctida Plate or Arctic Alaska/Chukotka Block
(Jones et al., 1987; Zonenshain and Natapov, 1989; Zonenshain et al., 1990) that includes the Brooks Range and North
Slope of northern Alaska. Neodymium isotope compositions
from the Brooks Range (Nelson et al., 1993) are the only
available data with which to characterize these crustal regions (Table 3); these data, and several exposures of Proterozoic rocks, suggest that the underlying basement of the
Arctic Alaska/Chukotka Block is in part Proterozoic in age,
as opposed to the neighboring continents (Le., Siberia and
North America), which have Archean cores.
It is widely agreed that at - 1 3 0 Ma the Canada Basin of
the Arctic Ocean opened by separation of the Arctic Alaska/
Chukotka Block from the Queen Elizabeth Island (QEI)
region, either by counterclockwise rotation or large scale
translational movement (cf. Grantz et al., 1979; Lawver and
Scotese, 1990; Moore et al., 1994). Chukotka, northern
Alaska, and the QEI (i.e., Innuition) regions share broadly
similar features, including elements of the Early Carboniferous Ellesmerian Orogeny, as well as lithologic and faunal
similarities (Trettin, 1989). However, it is probable that
clastics which comprise the thick sedimentary successions
of the QEI region (i.e., the Fanklinian Shelf and Sverdup
Basin) were mainly derived from the Archean North American craton to the south (Trettin, 1989), whereas clastic successions of the Alaska and Chukotka Foldbelts were derived
primarily from the Proterozoic Arctida Plate to the north
(i.e., prior to the opening of the Canada Basin; Zonenshain
4194
B. L. Winter et al.
Table 3. Ranges of present-day average neodymium and lead isotope compositions for potential sediment source
terranes to the central Arctic Ocean.
eNd(O)
PHANEROZOIC TERRANES:
Caledonian Orogenic Belts of Greenland, Scotland, Ireland, and Norway:
Caledonian Orogenic Belt and Shield Sediments of Baltic Shield:
Appalachian Orogenic Belt (on Phanerozoic basement):
Brooks Range, Alaska:
Northeast Siberia:
North-Central Siberia (inferred crustal contaminant of basalts):
PRECAMBRIAN TERRANES:
Northwest Canadian Shield (Wopmay, Rae, and Slave Provinces):
Mackenzie River Suspended Sediment
Archean of Labrador, Greenland, and Baltic:
Proterozoic of Greenland:
-8
-18
0
-2
to
to
to
to
?
- 1 0 to
-12
-22
-6
-11
-20
- 2 4 to - 3 0
-14.3
-45 to -25
-25 to -15
2°6PbflZ4pb Refs.
18.5-19.1
22-24
18.8-19.2
?
18.1-18.8
17.0-17.6
A
B
C
D
E
F
17-19
?
13-16
13-16
G
H
I
J
For Phanerozoid lead ore deposits, upper crustal estimates for 238U/2°4pb ratios (Stacey and Kramers, 1975) were
used to calculate present-day lead isotope compositions of the upper crust.
References:
A: Birkeland et al. (1993); Bjorlykke et al. (1993); Blaxland et al. (1979); Dixon et al. (1990); Hansen and
Friederichsen (1989); O'Nions et al. (1983); Sunblad and Stephens (1983).
B: Johansson and Rickard (1984); Miller et al. (1986); Romer and Wright (1993).
C: Ayuso and Bevier (1991); Barr and Hegner (1992); Foland and Allen (1991); Samson et al. (1995); Vitrac et
al. (1981); Wilbur et al. (1990).
D: Nelson et al. (1993).
E: Karpenko et al. (1981).
F: Lightfoot et al. (1993); Sharma et al. (1992); Wooden et al. (1993).
G: Davis and Hegner (1992); Gariepy and Allegre (1985); McCulloch and Wasserburg (1978); Miller et al. (1986);
Stevenson et al. (1987).
H: Goldstein et al. (1984).
I: Baadsgaard et al. (1986); Bennett et al. (1993); Cohen et al. (1991); Gariepy et al. (1990); Hamilton et al. (1978);
Hamilton et al. (1979); Hamilton et al. (1983); Kalsbeek et al. (1988); Kalsbeek et al. (1993); Moorbath et al.
(1986); Schiotte et al. (1986); Taylor et al. (1980); Taylor et al. (1992); Whitehouse (1989).
J: Kalsbeek et al. (1993); Patchett and Bridgwater (1984).
and Natapov, 1989; Moore et al., 1994). These relations
suggest that sediment derived from the QEI region will have
a greater ancient crustal component (i.e., low Er~dvalues and
high STSr/S6Sr ratios), which is probably best represented by
published crustal data from northwest Canada, including the
neodymium isotope datum of Mackenzie River sediment
(Table 3). In contrast, sediment derived from north Alaska
or the Chukotka Foldbelt of northeast Siberia is probably
best represented by published data from the Brooks Range
(Nelson et al., 1993). The shift to high 87Sr/86Sr ratios and
low eNa values during the last ~1.7 Ma in Arctic Ocean
sediments (Fig. 3) can, therefore, be explained by the QEI
region (Source 3) supplying a progressively greater proportion of the sediment over this time period.
Caledonian orogens that developed on or adjacent to Precambrian terranes that did not have widespread U-depletion
during granulite-facies metamorphism, such as the Baltic
Shield and parts of the Canadian Shield, tend to have very
radiogenic lead isotope ratios (Table 3; Fig. 7). Lead isotope variations for North Atlantic Ocean sediments clearly
indicate a major component from the Baltic Shield (Fig.
7), whereas this component is absent in sediment from the
Arctic Ocean. Lead isotope ratios for Arctic marine sediment are more similar to those of Caledonian orogenic
terranes that were built upon or adjacent to Archean crust
that was subject to ancient U-depletion, which produces
distinctly low 2°7pb/2°4pb rations as compared to average
upper crust (Fig. 7).
8. NEODYMIUM AND LEAD ISOTOPE COMPOSITIONS
OF ARCTIC SEAWATER
The neodymium and lead isotope compositions of Arctic
seawater during the Late Cenozoic, as determined from the
oxide fractions of Fe-Mn micronodules, closely track the
isotope compositions of the siliciclastic sediments. This correlation indicates that the ultimate sources for these metals
in Arctic seawater is similar to that of the siliciclastic sediments.
8.1. Neodymium Isotope Compositions
Present-day bottom waters of the central Arctic Ocean
have end values ranging from - 1 1 to - 1 3 (Fig. 3 and Table
4), slightly lower than those of Fe-Mn nodules from near
Svalbard (Cr~d = --9.7; Table 4) and seawater in the Greenland-Iceland-Norwegian Seas (ENd = --9.5 to --10.7; Table
4). The end values from the Arctic Ocean proper are, however, distinctly lower than those that were inferred for the
Arctic Ocean based on measurements of Arctic seawater
inflowing to Baffin Bay through Jones Sound (ENd = --9;
Stordal and Wasserburg, 1986). In comparison to the other
major ocean basins of the world, the central Arctic Ocean
is most similar to the Atlantic Ocean (eNd = --14 to --8;
Table 4), but significantly less radiogenic than the Indian
(eNd = --12 tO --5) and Pacific (eNd = --6 tO --3) Oceans.
The relatively low ENd values of the modern central Arctic
Ocean clearly indicate that older (i.e., Proterozoic-domi-
Provenance of Cenozoic sediment based on isotope ratios
Table 4. Neodymium isotope composition of present-day seawater.
Ocean Basin
end
Ref.
Atlantic
N. Atl. Deep Water
Indian
Pacific
Antarctic Bottom Water
Mediterranean Sea
Baffin Bay
Greenland-Iceland-NorwegianSeas
Denmark Strait Overflow
Jones Sound
Barents Sea
Amerasian Basin
- 16 to - 9
-13.5
- 11 to - 6
- 6 to - 2
-9
-7
-22 to - 19
- 10 to - 8
-9
-9
-10 to - 9
-13 to -11
1
2
1
1
3
4
5
2
2
5
6
this work
1.
2.
3.
4.
5.
6.
cf. Bertram and Elderfield, 1993.
Piepgras and Wasserburg, 1987.
Piepgras and Wasserburg, 1982.
Piepgras and Wasserburg, 1983.
Stordal and Wasserburg, 1983.
Amakawa et al., 1991.
eN d = (143Nd/144Ndsample- I43Nd/144NdcHuR)X 10 4.
nated) crustal terranes are important sources of Nd to Arctic
seawater.
8.2. Sources of Neodymium to Arctic Seawater
River water is the only well-documented source for the
dissolved rare earth element (REE) budget of seawater
(Goldstein and Jacobsen, 1987, 1988; Elderfield et al., 1990;
Sholkovitz, 1993), whereas, the contribution of REEs dissolved in seawater from particulate matter (aeolian, riverine,
hemipelagic, or ice rafted) via dissolution and/or exchange
processes is uncertain. Significant contributions of REEs to
seawater from aeolian particulates are supported by experimental studies (Greaves et al., 1994) and several seawater
investigations (North Atlantic: Elderfield and Greaves, 1982;
Mediterranean: Greaves et al., 1991; Thomas et al., 1994;
Henry et al., 1994; North Pacific: Shimizu et al., 1994). The
neoaymium isotope distribution pattern of detrital silicate
sediments in the Atlantic Ocean can be explained by mixing
between aeolian and fluvial inputs (Grousset et al., 1988),
and t:he similarity of this pattern to the geographic variations
of seawater neodymium isotope compositions suggests that
particulate matter may play an important role in controlling
the ~Nd value of Atlantic seawater (Grousset et al., 1988).
The much lower eNa value of Baffin Bay water (eNd = --20),
relative to that of the advected seawater input (eNd = --12
tO --'9), indicates that the surrounding Archean crust is contributing a significant amount of Nd to seawater within Baffin
Bay (Stordal and Wasserburg, 1986). The absence of fluvial
input: to Baffin Bay suggests that fine particulate material
transported by sea ice and glacial ice may be an important
source to the dissolved Nd fraction of seawater in ice-covered oceans (Stordal and Wasserburg, 1986). In contrast,
the large difference between the eNd values of central Pacific
seawater (eN~ = - 4 ) and that of the silicate sediment (eNd
= -- 10; Goldstein et al., 1984), which is exclusively aeolian,
very strongly suggests that volcanic ash, aeolian, and hemi-
4195
pelagic particles contribute a negligible proportion of REEs
to Pacific seawater (Jones et al., 1994).
Particulate matter rafted by sea ice and glacial ice is probably the major source of REEs to seawater in ice-covered
oceans such as the Arctic. It is improbable that changes in
the flux or isotope composition of Nd advected by external
seawater masses into the Arctic would produce simultaneous
changes in the eNd values of bottom waters and silicate sediment (Fig. 3). Furthermore, because all of the sediment we
have analyzed from the Alpha Ridge is ice rafted, it is unlikely that river water was the primary source of REEs to
Arctic seawater during the Late Cenozoic. For example, during the time period of sea ice sedimentation ( - 5 - 2 . 4 Ma),
river water could have only been the exclusive source of Nd
to seawater if the eNd value of the total river water input to
the Arctic was equal to that of the shelf sediment that was
delivered to the central Arctic Ocean by sea ice. This scenario is very improbable since the amount of sea ice sediment
delivered to the Arctic Ocean from a particular shelf region
is not dependent upon the local river discharge. Moreover,
because there are no rivers that drain into the Arctic Ocean
in the QEI region, it is unlikely that dissolved Nd supplied
by river water to the Arctic Ocean could have had the same
end value as the IRD during the time period of glacial sedimentation ( ~ 1.7-0 Ma). Therefore, variations in the neodymium isotope composition of Arctic seawater during the
Late Cenozoic reflect changes in the continental sources of
siliciclastic ice rafted debris.
8.3. Lead Isotope Compositions of Arctic Seawater
Similar to our arguments for Nd, we interpret that in general the covariation of lead isotope compositions of hydrogenous and siliciclastic sediment indicates that Pb in Arctic
seawater is predominantly derived from glacial and sea ice
rafted detritus (IRD) via exchange or partial dissolution processes. As with Nd, variations in the lead isotope compositions of Arctic seawater during the Late Cenozoic, therefore,
reflect the changes in continental sources of siliciclastic IRD.
For ocean basins other than the Arctic, Pb-Pb isotope variations appear to indicate mixing involving one endmember
that is similar to average crust (2°6pb/2°4pb = 18.7; 2°7pb/
2°4Pb = 15.63; 2°spb/2°4pb = 38.63, Stacey and Kxamers,
1975; Fig. 8). The lead isotope composition of the second
endmember for the Pacific data arrays is markedly less radiogenic than average crust, particularly in 2°Tpb/2°4Pb ratios
(Fig. 8), which suggests input of mantle-derived Pb from
either mid-ocean ridge hydrothermal processes or circumPacific undifferentiated arcs. The second endmember components for the Atlantic, Antarctic, and Indian mixing arrays
have lead isotope compositions that are substantially more
radiogenic than those of average crust (Fig. 8), which suggests input of Ph that was derived from distinctly different
ancient, upper crustal source terranes for each ocean basin.
In contrast to the other ocean basins of the world, neither
endmember of the lead isotope mixing arrays for the Arctic
Ocean is defined by average crustal Pb. The radiogenic component of the Arctic lead isotope mixing arrays reflects
Source 3 (Figs. 4 and 8), which is interpreted to be a Proterozoic-dominated, supracrustal source terrane in the northern
4196
B.L. Winter et al.
39.45
,
i
,
39.25
39.05
~ 38.85
O.
O 38.65
Ol
38.45
38.25
15.75
i
,
I
,
Antarctic-Atl.~
s-K Average
n
Antarctic-Pac.
\
Pscific
Antarctic-Ind.
2L
15.70
15.65
Anta " - .
Atlanticrct/~a
Antarctic-A~
15.60
c,,o
15.55
18.4
18.6
18.8
19.0
19.2
206pb1204pb
Fig. 8. Lead isotope compositions of the hydrogenous fraction of
Late Cenozoic Fe-Mn micronodules from the central Arctic Ocean.
Also shown are the data fields for Fe-Mn nodules from other major
ocean basins (Reynolds and Dasch, 1971; O'Nions et al., 1978; Ben
Othman et al., 1989; Abouchami and Goldstein, 1995). Stacey and
Kramers (1975) model for the lead isotope evolution of average
crust is shown for reference.
Canada or Queen Elizabeth Island region (see above).
Source 3 is distinct from the radiogenic Pb endmembers of
the mixing arrays defined by sediment from the other ocean
basins, in that the Arctic radiogenic component has relatively
low 2°spb//°apb ratios, reflecting time-integrated T h / U ratios
that are lower than those of average upper crust. The nonradiogenic component of the Arctic lead isotope mixing arrays
reflects Sources 1 and 2 (Figs. 4 and 8). The nonradiogenic
lead isotope compositions indicate that Sources 1 and 2 (Fig.
4) have comparatively lower time-integrated U/Pb, Th/Pb,
and U / T h ratios. We stress that the very low 2°Tpb/2°gPb and
2°6pb/E°4pb ratios, below those of average crust (Fig. 8), are
only found in ancient (Proterozoic or Late Archean) highgrade terranes. Because of the absence of circum-Arctic Tertiary arcs and the lack of ocean ridge hydrothermal activity
in the Amerasian Basin (Johnson, 1990), these lead isotope
compositions cannot reflect a direct mantle input, as is the
probable explanation for the low 2°7pb/2°4pb ratios of the
Pacific (Fig. 8).
9. CONCLUSIONS
Systematic variations in strontium, neodymium, and lead
isotope compositions of the silicate fractions of sediment
deposited in the central Arctic Ocean over the last 5 Ma
strongly coincide with a change in the sediment transport
mechanism from sea ice to glacial ice, concomitant with
changes in sediment provenance. Isotope correlation diagrams define three endmember sediment source compositions, which themselves probably represent mixtures of sedi-
ment derived from several crustal source regions. Between
5 and 1.7 Ma, sediment in the central Arctic Ocean was
deposited by sea ice. The relatively invariant strontium and
lead isotope ratios, combined with an increase in end values
at 3 Ma, suggests two distinct source regions for the silicate
fractions of this sea ice sediment ( 5 - 1 . 7 Ma). Because of
the absence of data from potential Arctic source areas, we
cannot determine the proportion of sediment that was derived
from a particular region. However, on the basis of modern
surface circulation in the Arctic Ocean and relatively low
2°6pb/2°4pb ratios of the older (5-1.7 Ma) Arctic sediment,
it is likely that a significant component of this sea ice sediment was derived from the East Siberian Shelf. The progressive increase in 87Sr/86Sr and 2°6pb/2°4pb ratios and the decrease in end values of the silicate fractions of sediment
deposited from 1.7 Ma to the present-day correlates with a
progressive increase in coarse glacially ice-rafted detritus
deposited over this time period. The isotope variations are
consistent with a progressively greater proportion of sediment derived from the northern Canada or Queen Elizabeth
Island regions since - 1 . 7 Ma.
Strontium isotope compositions of the oxide fractions of
Fe-Mn micronodules reflect those of >97% seawater Sr.
These results indicate that our chemical separation technique
was successful in concentrating the hydrogenous fraction of
the sediment. The lead and neodymium isotope compositions
of bottom Arctic seawater over the past 5 Ma coincide with
the isotope variations of the silicate fractions. These relations
indicate that particulate matter rafted by sea ice and glacial
ice is a major source of REEs and Pb to the dissolved reservoir in seawater in ice-covered oceans like the Arctic; river
water was probably not the primary source of these metals.
The eNd values of bottom water from the modern central
Arctic Ocean are most similar to those of the Atlantic and
indicate that older (i.e., Proterozoic-dominated) crustal terranes are important sources of Nd to Arctic seawater.
Acknowledgments--This work was supported by NSF grants OPP9122741 (DLC and CMJ), OPP-9400254 (DLC and BLW). We
thank B. Beard and K. Barovich for laboratory assistance. We thank
S. L. Goldstein, L. A. Derry, and an anonymous reviewer, as well
as GCA editor C. R. German, for helpful comments.
Editorial handling." C. R. German
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