PALEOCEANOGRAPHY, VOL. 19, PA4023, doi:10.1029/2003PA000994, 2004
Lower export production during glacial periods in the equatorial
Pacific derived from (231Pa//230Th)xs,0 measurements in deep-sea
sediments
Sylvain Pichat,1,2,3,4 Kenneth W. W. Sims,5 Roger François,6 Jerry F. McManus,5
Susan Brown Leger,6 and Francis Albarède1
Received 27 November 2003; revised 4 September 2004; accepted 30 September 2004; published 16 December 2004.
[1] The (231Pa/230Th)xs,0 records obtained from two cores from the western (MD97-2138; 1250S, 146240E,
1900 m) and eastern (Ocean Drilling Program Leg 138 Site 849, 011.590N, 11031.180W, 3851 m) equatorial
Pacific display similar variability over the last 85,000 years, i.e., from isotopic stages 1 to 5a, with systematically
higher values during the Holocene, isotopic stage 3, and isotopic stage 5a, and lower values, approaching the
production rate ratio of the two isotopes (0.093), during the colder periods corresponding to isotopic stages 2
and 4. We have also measured the 230Th-normalized biogenic preserved and terrigenous fluxes, as well as major
and trace elements concentrations, in both cores. The (231Pa/230Th)xs,0 results combined with the changes in
preserved carbonate and opal fluxes at the eastern site indicate lower productivity in the eastern equatorial
Pacific during glacial periods. The (231Pa/230Th)xs,0 variations in the western equatorial Pacific also seem to be
controlled by productivity (carbonate and/or opal). The generally high (231Pa/230Th)xs,0 ratios (>0.093) of the
profile could be due to opal and/or MnO2 in the sinking particles. The profiles of (231Pa/230Th)xs,0 and 230Thnormalized fluxes indicate a decrease in exported carbonate, and possibly opal, during isotopic stages 2 and 4 in
MD97-2138. Using 230Th-normalized flux, we also show that sediments from the two cores were strongly
affected by sediment redistribution by bottom currents suggesting a control of mass accumulation rates by
sediment focusing variability.
INDEX TERMS: 4231 Oceanography: General: Equatorial oceanography; 4267 Oceanography:
General: Paleoceanography; 4825 Oceanography: Biological and Chemical: Geochemistry; 4860 Oceanography: Biological and
Chemical: Radioactivity and radioisotopes; 4863 Oceanography: Biological and Chemical: Sedimentation; KEYWORDS: (231Pa/230Th)xs,0,
export productivity, Pacific
Citation: Pichat, S., K. W. W. Sims, R. François, J. F. McManus, S. Brown Leger, and F. Albarède (2004), Lower export production
during glacial periods in the equatorial Pacific derived from (231Pa/230Th)xs,0 measurements in deep-sea sediments, Paleoceanography,
19, PA4023, doi:10.1029/2003PA000994.
1. Introduction
[2] Increasing evidence for significant sea surface cooling
in the tropical ocean during glacial periods [Rosenthal et al.,
2003; Visser et al., 2004, and references therein] has lead to
a resurgence of interest in the role played by the equatorial
Pacific in Quaternary climatic cycles [e.g., Cane and
Clement, 1999]. The equatorial Pacific is one of the most
important sources of water vapor to the atmosphere and heat
1
Laboratoire de Sciences de la Terre, Ecole Normale Supérieure de
Lyon, Lyon, France.
2
Also at Department of Geology and Geophysics, Woods Hole
Oceanographic Institution, Woods Hole, Massachusetts, USA.
3
Also at Department of Marine Chemistry and Geochemistry, Woods
Hole Oceanographic Institution, Woods Hole, Massachusetts, USA.
4
Now at University of Oxford, Department of Earth Sciences, Oxford,
UK.
5
Department of Geology and Geophysics, Woods Hole Oceanographic
Institution, Woods Hole, Massachusetts, USA.
6
Department of Marine Chemistry and Geochemistry, Woods Hole
Oceanographic Institution, Woods Hole, Massachusetts, USA.
Copyright 2004 by the American Geophysical Union.
0883-8305/04/2003PA000994$12.00
to higher latitudes. Its impact on global climate is underscored by the perturbations associated with El Nino/Southern Oscillation (ENSO) cycles. El Nino is initiated by an
eastward displacement toward the central Pacific of the zone
of warmest surface waters, called the western Pacific warm
pool, and a weakening of the associated center of atmospheric convection. This shift results in a lower zonal sea
surface temperature (SST) gradient that weakens the Trade
Winds, decreases the equatorial upwelling and the zonal tilt
of the thermocline. In addition, the displacement of the
atmospheric convection center results in a global change in
atmospheric circulation that synchronously affects climate
in widespread regions of the globe. The system then swings
back to the opposite phase (La Nina) with the warm pool
moving and contracting westward, resulting in stronger
Trade Winds, equatorial upwelling and zonal thermocline
tilt.
[3] Building on these observations, it has been suggested
that changes in incoming solar radiation controlled by
orbital forcing could affect directly both the mean SST
and the SST distribution in the equatorial Pacific, and
change the frequency and intensity of ENSO. This, in turn,
would affect global climate through atmospheric telecon-
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1 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
nection, as observed today during El Nino events. Changes
in SSTs in the tropical Pacific could thus be one of the main
drivers of climatic variations during the late Quaternary,
particularly in the precessional frequency band [Cane, 1998;
Clement et al., 1999]. Only a few studies have shown a
precession-related signal in sediments from the equatorial
Pacific [Beaufort et al., 2001; Koutavas et al., 2002; Pichat
et al., 2003] therefore the significance of this forcing has
still to be established. Past ENSO variability may not
capture the full range of tropical climate variability on
Milankovich timescales, and whether glacial periods were
times of enhanced El-Nino or La Nina, or whether this
analogy is even appropriate are still open questions.
[4] With the modern orbital configuration, direct observations indicate that El Nino warms the North American
continent, while La Nina cools it [Cane, 1998]. It could thus
be surmised that, if similar atmospheric teleconnections
prevailed in the past, more frequent El Nino events would
have promoted melting of the Laurentide ice sheet, while
less frequent El Nino (or more prominent La Nina events)
would have promoted ice buildup [Cane and Clement,
1999]. Predominance of La Nina conditions during glacial
periods would be consistent with stronger Trade Winds,
during these periods, as evidenced by aeolian dust grain size
distribution [Parkin and Shackleton, 1973; Sarnthein et al.,
1981] and numerical simulations [e.g., Bush and Philander,
1999]. Belying this simple inference, however, are model
studies suggesting that solar forcing in the equatorial Pacific
would instead promote El Nino-like conditions during
glacial periods [Clement et al., 1999].
[5] Paleoceanographic evidence supporting either of these
scenarios is still ambiguous. It is now recognized that SST
in the equatorial Pacific was several degrees cooler during
the last glacial maximum [Lea et al., 2000; Kienast et al.,
2001; Stott et al., 2002; Rosenthal et al., 2003; Visser et al.,
2004]. While earlier reports of lower glacial SST in the
eastern equatorial Pacific (EEP) have been interpreted as
reflecting higher equatorial upwelling rates, i.e., a La Ninalike system, it has also been recognized that cooling could
arise from advection of cold water from the south [Lyle et
al., 1992; Mix et al., 1999; Feldberg and Mix, 2003] or
extratropical forcing [Andreasen et al., 2001]. Past changes
in the strength of the Trade Winds and equatorial upwelling
could be more confidently established from the zonal and
latitudinal SST gradients in the equatorial Pacific. Lea et al.
[2000] found a 3C drop in SST both in the eastern and
western Pacific, and a slightly larger zonal gradient during
glacial periods. They also found evidence for lower salinity
over the Ontong Java Plateau relative to the global ocean
mean. Both these observations are attributed to a La Ninalike situation, which is also consistent with the steeper zonal
tilt of the equatorial Pacific thermocline suggested by the
data of Andreasen and Ravelo [1997]. However, Koutavas
et al. [2002] found a smaller drop in glacial SST (1C) at a
site further south in the EEP. When compared to the results
obtained by Lea et al. [2000] further north and Kienast et al.
[2001] in the South China Sea, Koutavas et al. [2002]
results suggests weaker latitudinal and longitudinal SST
gradients in the glacial equatorial Pacific and weaker
upwelling, i.e., a dominance of El Nino. This interpretation,
PA4023
in turn, is consistent with the results of Stott et al. [2002]
who found saltier surface waters in the Mindanao Sea,
possibly reflecting the westward displacement of the center
of atmospheric convection, which characteristically happens
during El Nino events.
[6] The modern ENSO cycle affects primary production
in the equatorial Pacific [Murray et al., 1994]. Reconstructing past changes in equatorial productivity could
thus help establishing whether one ENSO mode (El Nino
or La Nina) prevailed in the past, or whether neither of
these climatic modes adequately describes the oceanography and atmospheric interactions of the glacial Earth.
Because El Nino curtails equatorial upwelling, lower
equatorial productivity is expected during periods when
it becomes prominent. Accurate estimation of past
changes in productivity is essential to fully describe and
understand the oceanography of the equatorial Pacific
during glacial periods and its impact on global climate.
In addition to being a diagnostic help, the evolution of
productivity in the tropical Pacific may have a direct
impact on global climate by affecting atmospheric CO2.
Changes in productivity and nutrient supply rate in the
equatorial Pacific would have little direct effect on the
atmospheric CO2, because all the nutrients upwelled at
the equator are eventually utilized by phytoplankton in
surface waters. However, changes in the ratio of silicate to
nitrate supply, as suggested by Matsumoto et al. [2002]
and Brzezinski et al. [2002], could significantly affect
atmospheric CO2 by altering plankton assemblages, organic
carbon to carbonate rain ratio and surface alkalinity.
[7] The importance of accurately reconstructing paleoproductivity in the equatorial Pacific has long been recognized
and has resulted in sustained efforts over the last decades.
Despite these efforts, a consensus has not yet been reached,
even as to whether productivity was higher or lower during
glacial periods. The more generally accepted view is that
glacial productivity in the EEP was higher, supporting a La
Nina-dominated glacial climate with stronger Trade Winds
and higher upwelling rates. This inference is mainly based
on accumulation rates of biogenic materials in the sediments
of the equatorial Pacific [Pedersen, 1983; Pedersen et al.,
1991; Lyle et al., 1988; Sarnthein et al., 1988; Paytan et al.,
1996]. However, this view has recently been challenged.
Variations in accumulation rates in the sediments of the
equatorial Pacific appear to be primarily driven by sediment
redistribution by bottom currents [Marcantonio et al., 2001;
Loubere et al., 2004]. Instead, a new array of paleoproductivity tracers points to lower glacial productivity in the
equatorial Pacific. For example, using a transfer function
based on benthic foraminifera assemblages, Loubere [1999,
2000, 2001, 2003] reports lower glacial productivity in the
South Equatorial Current (SEC) region of the EEP
(Figure 1), which is supplied with nutrients from the deeper
part of the equatorial undercurrent (EUC). In contrast, this
new tracer confirms higher glacial productivity in Panama
Basin and at the southern edge of the SEC region. Lower
glacial productivity in the SEC is also consistent with lower
carbonate rain rates obtained by combining 230Th-normalized sedimentary carbonate fluxes and estimates of carbonate preservation [Loubere et al., 2004].
2 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
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Figure 1. Cores location (black dots): ODP849 (Leg 138 Site 849, 011.590N, 11031.180W, 3851 m
water depth) in the eastern equatorial Pacific, MD2138 (1250S, 146240E, 1900 m water depth) in the
western part of the western Pacific warm pool and BC36 (0.14S, 158570E, 2311 m water depth) on the
Ontong Java Plateau. The main currents are also represented: north equatorial current (NEC), equatorial
undercurrent (EUC), south equatorial current (SEC), Peru current (PC), Mindanao Dome (MD) and New
Guinea coastal undercurrent (NGCU) originating from the SEC. PNG indicates the island of Papua New
Guinea.
[8] To further examine glacial/interglacial productivity
variability in the equatorial Pacific, we have measured
231
Pa, 230Th, 232Th, 235U, 238U, major and trace element
concentrations at three sites (Figure 1). We have calculated
the 230Th-normalized fluxes of biogenic (carbonate and
opal) and terrigenous material and the 231Pa to 230Th excess
activity ratio, decay corrected to the time of deposition
(hereafter referred to as (231Pa/230Th)xs,0). The latter has
been proposed as a proxy to assess changes in biological
productivity of the ocean during the last 150– 200 kyr [Lao
et al., 1992; François et al., 1993, 1997; Kumar et al., 1993,
1995], even though its interpretation can sometimes be
equivocal [e.g., Walter et al., 1999; Chase et al., 2002,
2003a]. The advantage of the (231Pa/230Th)xs,0 proxy is its
insensitivity to remineralization, dilution and sediment
redistribution. As yet, only a few studies [Lao et al.,
1992; Stephens and Kadko, 1997; Berelson et al., 1997]
have used this proxy to investigate productivity variations
during the Quaternary in the Pacific.
al., 1983a, 1983b; Bacon, 1984]. As a result, the flux of
scavenged 230Th to the seafloor approximates its known
production rate in the water column [Bacon, 1984;
Henderson et al., 1999; Yu et al., 2001] and can be used
as a reference to estimate sedimentary fluxes [Suman and
Bacon, 1989; François et al., 1990; McManus et al., 1998;
Frank et al., 1999; Chase et al., 2003b; François et al.,
2004].
[10] The total normalized flux or ‘‘rain rate’’ (RR) is
given by:
RR ¼ 230
bz
;
Thxs;0
ð1Þ
and the normalized flux for a component i (RRi), is given by
bzfi
;
RRi ¼ 230
Thxs;0
ð2Þ
where b is the constant production rate of 230Th from
U in the water column, b = 2.63 dpm/cm2/ka per km
of water depth; z is water depth, km; and fi is the weight
fraction of sedimentary constituent i. (230Thxs,0) is the
activity of scavenged 230Th corrected to the time of
deposition, dpm/gsediment.
[11] Because of its strong adsorption, scavenged 230Th
remains incorporated in sediment even if the particles that
originally transported it to the seafloor are solubilized
during early diagenesis. The ‘‘rain rates’’ calculated by
normalizing to 230Thxs,0 are thus ‘‘preserved’’ vertical
fluxes, i.e., the vertical fluxes of material that reach the
seafloor and remain after diagenetic remineralization.
[12] The focusing factor (y) is the ratio of the inventory
of 230Thxs,0 between dated horizons and 230Th produced in
seawater over the corresponding time interval [Suman and
Bacon, 1989]:
234
2. Principles Underlying the Use of (230Th)xs,0
Normalization and (231Pa/230Th)xs,0 to Estimate
Sedimentary Fluxes and Paleoproductivity
[9] Uranium is homogeneously distributed in the ocean
because of its long residence time (200– 450 kyr [Brewer,
1975; Ku et al., 1977; Chen et al., 1986]), compared with
the mixing time of the ocean (1000 – 1600 years [Broecker
and Peng, 1982]). As a consequence, 231Pa (t1/2 = 32 760
years) and 230Th (t1/2 = 75 380 years) are uniformly
produced in the water column at a constant activity ratio
of 0.093 (hereafter referred to as production rate ratio) from
a decay of 234U and 235U. Both 230Th and 231Pa are
extremely particle reactive, which leads to short residence
times in the water column: 10– 40 years for 230Th [Brewer
et al., 1980; Nozaki et al., 1981; Huh and Beasley, 1987]
and 50– 200 years for 231Pa [Anderson et al., 1983b; Nozaki
and Nakanishi, 1985; Yu et al., 1996]. With its higher
particle reactivity and shorter residence time, 230Th is
almost totally scavenged from the water column by the
vertical particle flux [Bacon and Rosholt, 1982; Anderson et
3 of 21
y¼
Z
r2
r1
230
Thxs;0 rr dr
bzðt1 t2 Þ
;
ð3Þ
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
where ri is sediment depth, cm; ti is the corresponding age
deduced from an independent chronology, ka; z is water
depth, km; and rr is dry bulk density, g/cm3.
[13] (230Thxs,0) and rr are averaged between dated sediment horizons ri. y > 1 indicates that more 230Thxs,0 has
accumulated than produced in the overlying water column,
thus indicating a lateral import of sediment to the area, i.e., a
net sediment focusing, whereas y < 1 indicates a net
winnowing.
[14] 231Pa has a lower particle affinity than 230Th and a
longer residence time in the water column. As a result,
231
Pa is more effectively transported over oceanic basinscale distances to be preferentially removed in areas of
higher particle flux and higher scavenging intensity
[Anderson et al., 1983b; Bacon, 1988; Yu et al., 2001].
This preferential removal, called boundary scavenging,
results in (231Pa/230Th)xs,0 > 0.093 in the sediments
underlying regions with high particle flux, which in open
ocean settings, often reflect higher export production [Yu et
al., 2001]. If particle flux were the only factor controlling
(231Pa/230Th)xs,0 in deep-sea sediments, this tracer could be
used as an unambiguous tracer of particle flux from which
paleoproductivity could be inferred. Ambiguities arise,
however, because (231Pa/230Th)xs,0 is also affected by
two other factors: (1) deep water circulation, which affects
the lateral transport of 231Pa within and between ocean
basins, and (2) particle composition, which influences the
relative affinity of 230Th and 231Pa for particles. The effect
of deep water circulation is best exemplified by contrasting
the increase in (231Pa/230Th)xs,0 with particle flux in the
Atlantic and Pacific ocean. Sediment trap experiments
have shown that the (231Pa/230Th)xs ratio is less sensitive
to particle flux in the Atlantic than in the Pacific [Yu et al.,
2001; Moran et al., 2002]. This difference has been
attributed to the shorter residence time of deep water in
the Atlantic (100 years [Broecker, 1979] versus 600 years
in the Pacific [Stuiver et al., 1983]) which limits the
establishment of lateral concentration gradients and prevents the full expression of boundary scavenging in the
Atlantic [Yu et al., 1996]. Therefore a single relationship
between (231Pa/230Th)xs,0 and particle flux, valid for all the
oceans, cannot be expected. Changes in particle composition further complicates the interpretation of sedimentary
(231Pa/230Th)xs,0. The (231Pa/230Th)xs,0 of sediment and
settling particles depends in part on their ferromanganese
oxides [Kadko, 1980; Anderson et al., 1983b; Shimmield
et al., 1986; Shimmield and Price, 1988; Frank et al.,
1994] and biogenic opal [Walter et al., 1997, 1999, 2001;
Chase et al., 2002, 2003a] content. Unlike the other
constituents of marine particles, ferromanganese oxides
have similar affinity for 231Pa and 230Th and biogenic
opal has a greater affinity for 231Pa than for 230Th. As a
result, their increasing prominence in settling particles
increases (231Pa/230Th)xs,0 in the underlying sediment,
independently of the particle flux. The effect of ferromanganese oxides is largely restricted to metalliferous sediments associated with hydrothermal plumes [Kadko, 1980;
Shimmield and Price, 1988] or the recycling of reduced
Mn from suboxic sediments [Shimmield and Price, 1988;
Anderson et al., 1983b]. Discerning the effect of opal is
PA4023
more difficult because of poor and variable preservation.
Thus opal concentration in sediment does not necessarily
reflect opal content in settling particles that brought 230Th
and 231Pa to the seafloor. Using sediment trap samples
ranging from opal-dominated to carbonate-dominated
regions, Chase et al. [2002] found a strong correlation
between the opal/carbonate ratio and the (231Pa/230Th)xs of
settling particles. However, within the equatorial Pacific
region, where the variability of opal/carbonate ratio is
smaller, (231Pa/230Th)xs correlates better with particle flux
than with particle composition. As a result, we are not able
as yet to distinguish clearly between the relative importance of particle flux and particle composition in controlling (231Pa/230Th)xs,0. Therefore observed past changes in
sedimentary (231Pa/230Th)xs,0 in the equatorial Pacific
could reflect either changes in particle flux, or in the
relative opal content of settling particles or both. To better
constrain our observations, we have also measured the
total, biogenic (opal and carbonate) and terrigenous 230Thnormalized fluxes, as well as concentrations of major and
trace elements in the sediments.
3. Experimental Section
3.1. Sediment Samples
[15] Three cores located along the Equator have been
analyzed in this study (Figure 1). Core MD97-2138 (hereafter referred to as MD2138) was collected using the
CALYPSO Kullemberg giant piston corer aboard the R/V
Marion Dufresne during campaign IMAGES III. The
MD2138 site is situated at a depth of 1900 m, north of
Manus Island, 300 km north of Papua New Guinea, in the
western part of the Western Pacific Warm Pool. ODP Leg
138 Site 849 cores (hereafter referred to as ODP849) are
located in the eastern equatorial Pacific (EEP) in deeper
water (3800 m) at about 850 km west of the East Pacific
Rise. Core MW91-9 BC36, referred hereafter to as BC36,
has been collected using a box corer during R/V Moana
Wave cruise 9 on the Ontong Java plateau (2311 m water
depth).
[16] All three cores are located above the carbonate
compensation depth and far from any hydrothermal sources.
ODP849 and BC36 are far from riverine sources of terrigenous material. For these two cores, particulate matter
sinking through the water column is predominantly biogenic
and winds are the only significant supplier of terrigenous
material. On the contrary, MD2138 is located close to Papua
New Guinea where large riverine discharges of terrigenous
material occur [Milliman et al., 1999].
3.2. Age Models and Stratigraphy
[17] For ODP849, we used the age model of Mix et al.
[1995] based on the comparison between d18O record on
benthic foraminifer C. wuellerstorfi and the SPECMAP
stack [Imbrie et al., 1984]. For core MD2138, we used an
age model based on six 14C ages and the d18O record of
planktonic foraminifer G. ruber (T. de Garidel-Thoron,
manuscript in preparation, 2004). For core BC36, we
have used the d18O records on planktonic foraminifer,
P. obliquiloculata and G. sacculifer, from Patrick and
4 of 21
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Table 1. Chemical Data and Estimates of Percent Opal, Percent Carbonate, and Percent MnO2 Obtained by Normative Calculationsa
Sample
Depth, cm
Al, wt %
Ca, wt %
Fe, wt %
Mn, wt %
Si, wt %
Ti, wt %
Opal, %
CaCO3, %
Lithog, %
MnO2, %
MD01
MD02
MD08
MD13
MD14
MD19
MD21
MD23
MD26
MD34
MD38
MD46
MD48-a
MD48-b
MD48-c
MD48-ave
MD52
MD53
MD55
MD58
MD60
MD64
2s, %
3.5
4.5
10.6
14.7
15.5
19.4
21.0
22.8
25.7
35.7
40.8
51.9
54.7
54.7
54.7
54.7
60.9
62.6
65.9
71.3
75.0
80.9
3.25
12.25
72.25
122.25
132.25
182.25
202.25
222.25
252.25
332.25
372.25
402.25
472.25
472.25
472.25
472.25
502.25
522.25
542.25
572.25
582.25
622.125
3.49
2.71
2.25
2.56
2.84
3.48
3.26
3.71
3.05
2.47
2.17
2.40
2.10
2.29
2.24
2.21
2.41
2.57
2.64
3.33
3.50
2.34
2.9
25.0
22.4
27.5
23.4
22.7
20.9
19.0
19.5
21.6
31.9
23.6
25.1
25.4
26.8
27.2
26.5
24.6
25.5
21.2
22.6
22.3
24.9
4.0
2.47
1.64
1.65
2.08
2.07
2.55
2.31
2.99
2.21
1.83
1.44
1.56
1.61
1.58
1.58
1.59
1.69
1.80
2.09
1.99
2.77
1.74
4.9
MD2138
0.38
0.15
0.03
0.03
0.03
0.05
0.04
0.05
0.04
0.04
0.03
0.04
0.03
0.04
0.04
0.04
0.04
0.05
0.04
0.05
0.05
0.04
4.3
11.7
9.6
6.5
8.0
8.5
11.0
9.8
11.9
9.3
7.3
6.4
7.2
6.1
6.8
6.5
6.5
7.0
7.8
7.9
9.5
12.0
6.8
6.8
0.18
0.13
0.12
0.14
0.15
0.19
0.19
0.21
0.17
0.13
0.12
0.13
0.11
0.12
0.12
0.12
0.13
0.14
0.14
0.18
0.19
0.12
4.0
3.6 – 5.1
3.9 – 5.1
0 – 0.9
1.1 – 2.2
0 – 1.5
1.9 – 3.4
0 – 1.9
2.5 – 4.1
0 – 2.2
0 – 1.2
0 – 1.0
0 – 1.4
0 – 0.9
0 – 1.2
0 – 0.8
0 – 0.9
0 – 0.9
0 – 1.8
0 – 1.4
0 – 0.7
4.2 – 5.8
0 – 0.8
55.6 – 61.6
50.6 – 55.2
64.4 – 68.3
53.5 – 57.9
51.2 – 56.1
45.3 – 51.3
41.2 – 46.8
41.5 – 47.8
47.9 – 53.2
74.8 – 79.0
54.7 – 58.4
58.0 – 62.1
59.3 – 62.9
62.4 – 66.4
63.6 – 67.4
61.8 – 65.6
56.8 – 61.0
58.6 – 63.0
47.8 – 52.4
50.0 – 55.7
48.8 – 54.8
57.6 – 61.6
34.9 – 38.8
27.1 – 30.1
22.5 – 25.0
25.6 – 28.5
28.4 – 31.6
34.8 – 38.7
32.6 – 36.2
37.1 – 41.2
30.5 – 33.9
24.7 – 27.5
21.7 – 24.1
24.0 – 26.7
21.0 – 23.3
22.9 – 25.5
22.4 – 24.9
22.1 – 24.6
24.1 – 26.8
25.7 – 28.5
26.4 – 29.3
33.3 – 37.0
35.0 – 38.9
23.4 – 26.0
0.51 – 0.55
0.17 – 0.20
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
OPD032
ODP072-a
ODP072-b
ODP072-ave
ODP082
ODP102
ODP122-a
ODP122-b
ODP122-ave
ODP182
ODP202
ODP222
ODP232
ODP272
2s, %
8.25
17.07
17.07
17.07
18.95
23.57
28.63
28.63
28.63
48.57
55.79
63.19
66.37
78.87
31.8
73.7
73.7
73.7
82.3
102.1
121.8
121.8
121.8
182.1
201.7
222.3
231.9
272.4
0.32
0.24
0.26
0.25
0.20
0.16
0.15
0.11
0.13
0.14
0.40
0.66
0.45
0.51
1.1
28.8
31.2
31.0
31.1
30.7
31.2
26.5
31.4
29.0
31.2
25.8
25.3
24.3
23.2
3.4
0.29
0.20
0.20
0.20
0.19
0.18
0.15
0.16
0.16
0.19
0.36
0.51
0.38
0.38
4.2
ODP849
0.34
0.12
0.13
0.13
0.18
0.13
0.10
0.10
0.10
0.16
0.84
0.29
0.15
0.42
3.6
6.9
3.7
3.8
3.8
3.3
3.8
4.3
4.3
4.3
6.1
9.1
9.5
8.4
10.3
4.5
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
5.1
13.5 – 14.2
6.7 – 7.2
6.7 – 7.3
6.7 – 7.2
6.1 – 6.5
7.6 – 7.9
8.9 – 9.3
9.2 – 9.4
9.1 – 9.4
13.4 – 13.7
18.2 – 19.1
16.9 – 18.3
16.0 – 16.9
20.2 – 21.3
71.5 – 71.9
77.5 – 77.9
77.1 – 77.5
77.3 – 77.7
76.3 – 76.6
77.7 – 77.9
66.0 – 66.2
78.3 – 78.4
72.2 – 72.3
77.6 – 77.8
63.9 – 64.4
62.2 – 63.1
59.9 – 60.5
57.0 – 57.7
3.2 – 3.9
2.4 – 3.0
2.6 – 3.2
2.5 – 3.1
2.0 – 2.5
1.6 – 1.9
1.5 – 1.9
1.1 – 1.4
1.3 – 1.6
1.4 – 1.8
4.0 – 5.0
6.6 – 8.2
4.5 – 5.6
5.1 – 6.3
0.53
0.19
0.20
0.19
0.27
0.20
0.15
0.16
0.15
0.25
1.32
0.44
0.23
0.65
PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
5 of 21
Age, ka
a
Percent lithogenic was obtained by difference in MD2138 and from percent Al in ODP849 (see text). Full replicate analyses are indicated by a dash followed by a letter after the name of the sample, and the
average of the replicate measurements is indicated by ‘‘-ave’’ after the name of the sample. BDL, below detection limit. Precision (2 s) is reported based on 5 replicate measurements of a standard (USGS marine
core MAG-1) for MD2138, respectively 7 for ODP849.
PA4023
PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
PA4023
PA4023
Table 2. Isotopic Data for Cores MD2138, BC36, and ODP849a
Depth, cm
(238U), dpm/g
3.5
4.5
4.5
5.7
5.7
6.8
8.8
10.6
10.6
10.6
10.6
10.6
12.3
14.0
14.7
14.7
15.5
16.3
17.1
17.8
18.6
19.4
20.2
20.2
20.2
20.2
21.0
21.9
22.8
23.7
24.7
25.7
26.8
28.0
29.3
32.0
32.0
32.0
35.7
40.8
40.8
48.9
48.9
48.9
48.9
48.9
51.9
53.2
54.7
59.3
60.9
62.6
64.2
65.9
67.7
69.4
69.4
71.3
75.0
80.9
80.9
3.25
12.25
12.25
22.25
22.25
32.25
52.25
72.25
72.25
72.25
72.25
72.25
92.25
112.25
122.25
122.25
132.25
142.25
153.25
162.25
172.25
182.25
192.25
192.25
192.25
192.25
202.25
212.25
222.25
232.25
242.25
252.25
262.25
272.25
282.25
303.25
303.25
303.25
332.25
372.25
372.25
432.25
432.25
432.25
432.25
432.25
453.25
462.25
472.25
502.25
512.25
522.25
532.25
542.25
552.25
562.25
562.25
572.25
592.25
622.125
622.125
0.42
0.49
0.44
0.31
0.31
0.36
0.42
1.68
1.68
1.66
1.64
1.64
1.91
2.85
2.31
2.52
2.38
2.22
2.62
2.85
2.48
2.53
2.72
2.76
2.72
3.01
3.28
2.84
2.32
1.94
2.51
3.33
3.50
2.67
3.39
3.70
3.58
3.57
2.80
2.76
2.76
3.30
3.30
3.16
3.20
3.20
2.16
2.63
2.71
2.39
2.83
2.79
2.32
0.49
2.12
2.31
2.24
2.28
3.37
2.10
2.19
0.4
0.4
2.8
6.8
11.1
1.5
1.5
10.5
21.0
28.5
0.19
0.19
0.18
0.16
0.14
Sample
Age, ka
MD01
MD02-a
MD02-b
MD03-1
MD03-2
MD04
MD06
MD08-b
MD08-c1
MD08-c2
MD08-d1
MD08-d2
MD10
MD12
MD13-a
MD13-b
MD14
MD15
MD16
MD17
MD18
MD19
MD20-a
MD20-b
MD20-c
MD20-d
MD21
MD22
MD23
MD24
MD25
MD26
MD27
MD28
MD29
MD31-a
MD31-b
MD31-c
MD34
MD38
MD38
MD45-a1
MD45-a2
MD45-b
MD45-c1
MD45-c2
MD46
MD47
MD48
MD51
MD52
MD53
MD54
MD55
MD56
MD57-a
MD57-b
MD58
MD60
MD64-a
MD64-b
BC36#1-1
BC36#1-2
BC36#2
BC36#3
BC36#4
(232Th), dpm/g
(230Th)xs,0, dpm/g
(231Pa)xs,0, dpm/g
(231Pa/230Th)xs,0
MD2138
0.28
0.25
0.23
0.18
0.18
0.18
0.20
0.24
0.23
0.23
0.23
0.23
0.30
0.32
0.35
0.35
0.39
0.35
0.34
0.33
0.35
0.34
0.35
0.31
0.31
0.34
0.39
0.37
0.40
0.40
0.36
0.39
0.36
0.33
0.35
0.33
0.29
0.29
0.25
0.26
0.26
0.29
0.29
0.38
0.21
0.21
0.23
0.27
0.26
0.32
0.35
0.34
0.37
0.25
0.40
0.34
0.33
0.38
0.24
0.26
0.26
3.84
3.96
4.12
3.84
3.84
2.98
3.69
3.79
3.75
3.75
3.66
3.66
3.00
3.27
3.02
3.38
3.72
3.34
3.63
3.47
3.25
3.57
3.74
3.64
3.81
3.79
4.13
3.99
3.58
3.42
4.00
4.26
3.50
3.11
3.49
3.36
3.78
3.80
3.36
3.36
3.36
3.84
3.84
4.00
4.04
4.05
3.13
3.29
2.63
2.92
3.35
3.56
3.82
6.73
3.74
4.66
4.60
4.04
3.90
2.98
2.97
0.49
0.45
0.49
0.48
0.48
0.49
0.55
0.47
0.44
0.43
0.43
0.44
0.36
0.37
0.22
0.36
0.42
0.35
0.42
0.40
0.34
0.43
0.40
0.45
0.47
0.41
0.42
0.47
0.34
0.41
0.43
0.39
0.42
0.37
0.42
0.41
0.49
0.51
0.46
0.45
0.45
0.53
0.53
0.57
0.55
0.55
0.47
0.37
0.35
0.37
0.32
0.40
0.43
0.69
0.50
0.58
0.63
0.53
0.55
0.41
0.47
0.129
0.113
0.120
0.125
0.126
0.166
0.149
0.123
0.119
0.115
0.119
0.121
0.120
0.112
0.073
0.108
0.114
0.104
0.116
0.114
0.105
0.121
0.106
0.123
0.123
0.107
0.103
0.118
0.094
0.119
0.107
0.091
0.119
0.118
0.121
0.124
0.129
0.133
0.137
0.135
0.135
0.139
0.137
0.143
0.135
0.134
0.152
0.113
0.132
0.127
0.095
0.112
0.114
0.102
0.133
0.125
0.136
0.132
0.142
0.138
0.157
BC36
0.11
0.11
0.09
0.08
0.10
5.15
5.15
4.96
4.49
4.17
0.48
0.48
0.45
0.40
0.39
0.092
0.093
0.092
0.090
0.093
6 of 21
PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
PA4023
PA4023
Table 2. (continued)
Sample
ODP012
ODP032
ODP042
ODP052
ODP062
ODP072
ODP082
ODP102
ODP112
ODP122
ODP142
ODP162-a
ODP162-b
ODP162-c
ODP182
ODP202
ODP222
ODP232
ODP252
ODP272
ODP292
Age, ka
4.35
8.25
10.52
12.53
14.56
17.07
18.95
23.57
26.04
28.63
34.36
41.07
41.07
41.07
48.57
55.79
63.19
66.37
72.59
78.87
85.36
Depth, cm
(238U), dpm/g
11.8
31.8
42.2
51.8
61.7
73.7
82.3
102.1
112.1
121.8
141.6
162.1
162.1
162.1
182.1
201.7
222.3
231.9
251.9
272.4
291.8
0.27
0.19
0.19
0.17
0.13
0.12
0.12
0.12
0.11
0.15
0.12
0.12
0.12
0.12
0.12
0.16
0.18
0.14
0.12
0.16
0.13
(232Th), dpm/g
ODP849
0.03
0.03
0.05
0.06
0.05
0.07
0.08
0.06
0.05
0.04
0.05
0.05
0.03
0.03
0.05
0.10
0.14
0.12
0.06
0.09
0.10
(230Th)xs,0, dpm/g
(231Pa)xs,0, dpm/g
(231Pa/230Th)xs,0
11.11
8.95
9.25
9.96
10.52
9.91
11.10
9.00
8.50
7.91
8.31
9.07
9.05
8.93
7.55
14.53
21.72
20.07
13.40
18.75
20.39
1.29
1.06
1.08
1.10
1.02
0.91
0.99
0.86
0.71
0.79
0.89
0.92
0.89
0.91
0.84
1.28
1.86
1.84
1.45
1.72
1.99
0.116
0.119
0.116
0.110
0.097
0.092
0.089
0.095
0.084
0.100
0.107
0.101
0.098
0.102
0.111
0.088
0.086
0.092
0.108
0.092
0.098
a
Full replicates are indicated by a dash followed by a letter after the name of the sample. Replicate analyses are indicated by a dash followed by a number
after the name of the sample. Parentheses denote activity.
Thunell [1997] and two bulk sediment
et al., 1999].
14
C ages [Broecker
3.3. Analytical Procedures
3.3.1. Elemental analyses
[18] The concentrations (wt %) of Al, Ca, Fe, Mn, Si, and
Ti were measured in 19 samples of MD2138 and 10 samples
of ODP849 (Table 1). Analyses were performed by
inductively coupled plasma optical emission spectrometry
(ICP-OES, Jobin Yvon JY38VHR). Accuracy was within
3% as determined by repetitive measurements of a
standard (USGC marine sediment MAG-1). Precisions
are reported in Table 1. The carbonate, opal and lithogenic content of the sediments were estimated by normative calculations assuming that all the Al is associated
with the terrigenous fraction (see section 4.1) and using a
mean detrital material chemical composition to estimate
the terrigenous contribution for the other elements (i),
according to:
h
i
ðiÞbiogenic ¼ ðiÞtotal ¼ ði=AlÞj ðAlÞtotal ;
ð4Þ
where j is bulk continental crust (CC), upper continental
crust (UPCC), terrigenous post-Archean Australian terrigenous shales (PAAS) as defined by Taylor and McLennan
[1985] or mafic rocks from the Manus basin (MRMB) as
given in the text (see section 4.1).
3.3.2. The 231Pa and 230Th Analyses
[19] 231Pa and 230Th (Table 2) were measured by
isotopic dilution using 233Pa and 229Th spike, respectively, on a single collector, sector field inductively
coupled plasma mass spectrometer (SF-ICP-MS), following the procedures described in the work of Choi et al.
[2001], Pichat [2001] and S. Pichat et al. (manuscript in
preparation, 2004). Briefly, the sediment samples were
spiked and equilibrated with 233Pa and 229Th prior to total
dissolution in HNO3, HF and HClO4. An aliquot, representing less than 1 wt %, of the resulting solution was
analyzed directly for 238U and 232Th by isotope dilution
using 236U and 229Th spikes. The remaining solution was
used for 231Pa and 230Th separation by ion-exchange
chromatography, slightly modified from Fleer and Bacon
[1991] and subsequently analyzed by SF-ICP-MS (Finnigan, Element) in low resolution mode. The instrumental
mass fractionation was evaluated by bracketing each
sample measurement with analyses of an uranium standard (National Bureau of Standards NBS 960) (S. Pichat
et al., manuscript in preparation, 2004). The signal was
corrected from the contributions of the instrumental
background, dark noise, blanks linked to the chemical
procedure, and blanks linked to spike addition (S. Pichat
et al., manuscript in preparation, 2004). The sum of these
contributions was always <0.5%, and usually <0.1%, of
the signal. The internal precision, i.e., the calculated
uncertainty based on two times the standard error (2s)
propagated from the 231Pa, 233Pa, 230Th and 229Th measurements, on the (231Pa/230Th)xs,0 ratios was usually better
than 1.5%. Reproducibility for full replicate analyses
(dissolution, chromatographic separation and spectrometry) was usually better than 5% (2s) on the
(231 Pa/ 230 Th) xs,0 ratio. Initial excess activities were
obtained after corrections for (1) the detrital 230Th or
231
Pa contribution, estimated from the 232Th content of
the sediment and using the average (238U/232Th) activity
ratio of the lithogenic fraction of the marine sediments:
0.8 ± 0.2 [Anderson et al., 1990] for ODP849 and BC36,
1.0 ± 0.3 for MD2138 (see section 4.1), (2) the decay
since the time of sediment deposition estimated from the
age model of each core, and (3) the postdepositional
7 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
Figure 2. Elemental ratios measured in cores (left) MD2138 and (right) ODP849. The horizontal lines
represent: terrigenous post-Archean Australian terrigenous shales (PAAS, thick dashed line), upper
continental crust (UPCC, thick continuous line), bulk continental crust (CC, thick dot-dashed line)
[Taylor and McLennan, 1985], mafic rocks from the Manus basin (MRMB, thin continuous line, data
compiled from Stracke and Hegner [1998] and Gill et al. [1993], see section 4.1 for explanations) and
island arcs (thin dashed line) [Taylor and McLennan, 1995]. BDL, below detection limits.
8 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
PA4023
Figure 3. Sediment composition deduced from elemental
ratios (large symbols). The smaller connected symbols for
MD2138 represent estimates derived from our high
resolution (HR) 232Th measurements (see text for explanations). For MD2138, reported preserved opal represents the
maximum values, i.e., using a pure MRMB end-member
(MRMB: mafic rocks from the Manus basin, see section 4.1
for explanations) for the lithogenic fraction of the sediment.
Glacial periods (isotopic stages 2 and 4) are shown in light
gray [Martinson et al., 1987].
ingrowth of
present.
230
Th or
231
Pa when authigenic uranium was
4. Results
4.1. Elemental Composition of Sediments From Sites
MD2138 and ODP849
[ 20 ] The elemental compositions of MD2138 and
ODP849 (Table 1) show a clear contrast between the two
cores. Given the location of MD2138, the lithogenic fraction of the sediment is likely to be dominated by mafic
rocks. The 232Th/Ti and Ti/Al ratios were used to estimate
the chemical composition of the lithogenic components of
MD2138 sediment. The 232Th/Ti ratios of MD2138 lie
between two end-members: a low 232Th/Ti ratio represented
by mafic rocks, either MRMB (a compilation of mafic rocks
data from the Manus basin) or island arcs [Taylor and
McLennan, 1995], and a high 232Th/Ti ratio represented
either by the terrigenous PAAS (post-Archean Australian
terrigenous shales) or the UPCC (upper continental crust)
(Figure 2). The Ti/Al values of MD2138 (Figure 2) also lie
between those of the mafic rocks and the terrigenous PAAS
PA4023
or the UPCC. However, the 232Th/Ti panel shows that if the
UPCC contributes to the detrital fraction of MD2138 sediments, it is likely to be insignificant. The elemental ratios of
MRMB and island arcs are very similar (Figure 2). We have
chosen to use our compiled mafic rocks values from the
Manus basin (MRMB) [Gill et al., 1993; Stracke and
Hegner, 1998] to make the normative calculations rather
the average island arcs values of Taylor and McLennan
[1995] in order to match the geographical settings of
MD2138. We have therefore chosen MRMB and PAAS as
end-members for the detrital fraction of the MD2138
sediment. In MD2138, Ca/Al is much higher than expected
from MRMB or PAAS while Si/Al is close to lithogenic
material values. These elemental compositions clearly indicates that the sediment accumulated at the western Pacific
site (MD2138) is dominated by biogenic carbonate and
lithogenic material. Because of its high concentration, the
carbonate content can be accurately estimated by normative
calculations (equation (4)). There is no difference in using
Ti instead of Al as the normative element in MD2138 which
shows that there is probably no, or very little, Alxs [Murray
et al., 1993]. Thus we choose to use Al rather than Ti
because concentrations for the latter are low in MD2138 and
below detection limits for ODP849 (Table 1). Calculations
using CC, UPCC, PAAS or MRMB show an overall
variation of less than 6 wt % in the carbonate estimate
and less than 4 wt % in the lithogenic estimate. The range of
values given in Table 1 reflects the use of MRMB or PAAS
as lithogenic component for the normative calculations.
Percent biogenic Si is estimated by normative calculations
using a lithogenic Si/Al = 2.75 for the MRMB and 2.93 for
the PAAS [Taylor and McLennan, 1985]. A factor of 2.4 is
then used to convert percent biogenic Si to percent biogenic
opal [Mortlock and Froelich, 1989]. Calculations based on
the PAAS show no preserved opal except 3.5 wt % in the
top 15 cm of the core. An average value of 2 wt %
preserved opal is calculated when using the MRMB endmember. In this latter case, the upper 15 cm of MD2138
show a preserved opal of 5 wt % (Table 1). Overall, given
the uncertainties of the measurements, the preserved opal
fraction in MD2138 is likely to be insignificant. From the
analyses of the 19 MD2138 samples, we can deduce the
mean 232Th concentration of the lithogenic sediment deposited at this site (3.8 ± 0.6 ppm), and use the 232Th
concentrations measured at higher resolution (Table 2) to
derive higher resolution profiles of percent lithogenic and
percent carbonate (Figure 3). From the mean 232Th concentration, we can also calculate the relative contribution of the
two end-members, MRMB (232Th 1.15 ppm) and PAAS
(232Th 14.6 ppm), to the lithogenic fraction assuming a
binary mixing. We found that the MRMB contributes to 80 ±
5% of the lithogenic fraction of the sediment (Table 3). This
result allows us to calculate an average 238U over 232Th
activity ratio for the lithogenic fraction of MD2138:
(238U/232Th)litho = 1.0 ± 0.1 (Table 3). However, we choose
to use a wider range of (238U/232Th)litho = 1.0 ± 0.3 to take
into account the uncertainties in the estimation of the
relative contributions of the two end-members.
[21] In contrast to MD2138, elemental ratios at the
ODP849 site (Table 1; Figure 2) reveal the presence of
9 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
PA4023
Table 3. Estimation of the (238U/232Th) Activity Ratio of the Lithogenic Fraction of MD2138a
MRMB
PAAS
Island arcs
75% MRMB + 25% PAAS
80% MRMB + 20% PAAS
85% MRMB + 15% PAAS
Th, ppm
U, ppm
(Th), dpm/g
(U), dpm/g
(238U/232Th)
1.15
14.6
2.5
4.5
3.8
3.2
0.82
3.1
1.0
1.4
1.3
1.1
0.28
3.55
0.61
1.10
0.93
0.77
0.61
2.31
0.75
1.04
0.95
0.87
2.18
0.65
1.23
0.94
1.02
1.12
a
MRMB, mafic rocks from the Manus Basin (see text); PAAS, post-Archean Australian terrigenous shales [Taylor and McLennan,
1985]. The average island arcs values [Taylor and McLennan, 1995] are given for comparison. Parentheses denote activity.
significant biogenic opal burial in addition to carbonate.
Both Al and 232Th concentrations indicate that the lithogenic content is about ten-fold lower than at the western
site, and account for only a few percents of the sediment
mass (Table 1). Ti concentrations are below detection limits.
Under these circumstances, the dominant carbonate phase
can be accurately determined from percent Ca after a minor
correction for lithogenic contribution, using equation (4)
(Figure 3). The percent preserved opal is estimated as for
MD2138 (Figure 3). The small lithogenic component is best
quantified from Al concentrations (Figure 3). The range of
values given in Table 1 reflects the use of PAAS, UPCC or
CC as lithogenic component. Using PAAS, UPCC or CC for
the normative calculations results in variations lower than
0.6% for the different fractions.
[24] The results clearly confirm reducing conditions, i.e.,
high authigenic U activity, below 60 cm depth in core
MD2138. However, very little or no authigenic U accumulates in ODP849.
4.3. Variations in (231Pa//230Th)xs,0
[25] The short Ontong Java Plateau core (BC36) records
a nearly constant (231Pa/230Th)xs,0 over the entire Holocene that is very similar to the production rate ratio
(Figure 5). However, the two longer records from the
western (MD2138) and eastern (ODP849) equatorial
Pacific display systematic and similar variations in
(231Pa/230Th)xs,0, with minima close to the production
rate ratio during the colder isotopic stages 2 (IS2) and
4 (IS4), and maxima during the warmer isotopic stages 1
4.2. Contrasting Redox States Between Core MD2138
and ODP849
[22] In the western Pacific core, Mn/Al lie between the
MRMB and the PAAS values, except for the core top,
showing evidence of Mn diagenetic reductive remobilization (Figure 2). MD2138 sediment becomes sufficiently
reducing between 12 cm and 72 cm depth to reduce
MnO2. In contrast, Mn/Al ratio remains very high compared
to the CC, UPCC or PAAS values over the entire length of
ODP849, indicating that the sediment deposited at this site
during the last 80 ka remained oxic. MnO2 concentrations
(Table 1) were estimated by normative calculations using
PAAS or MRMB Mn/Al values for MD2138 and CC,
UPCC or PAAS [Taylor and McLennan, 1985] Mn/Al
values for ODP849. For the eastern Pacific, normative
calculations using the Mn/Al ratios of CC, UPCC or PAAS
give the same MnO2 concentrations.
[23] The contrast in redox conditions between the two
cores is confirmed by their respective authigenic U concentrations (Figure 4). Uranium diffuses from bottom water and
precipitates in suboxic and anoxic sediment pore waters
[e.g., Barnes and Cochran, 1990; Mangini et al., 2001].
This authigenic fraction of sedimentary U can be estimated
by normative calculations using the average (238U/232Th)
activity ratio of the lithogenic fraction of the marine sediments (0.8 ± 0.2 [Anderson et al., 1990] for ODP849 and
BC36; 1.0 ± 0.3 for MD2138):
h
238
U authigenic ¼ 238 U total 238 U=232 Th
litho
232 i
Th total :
ð5Þ
Figure 4. MnO2 and authigenic U concentrations in
MD2138 and ODP849 calculated by normalization to Al
for MnO2 and to 232Th for U. rcmcd: revised centimeters
composite depth [Mix et al., 1995].
10 of 21
PA4023
PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
PA4023
Figure 5. (231Pa/230Th)xs,0 and d18O in MD2138, BC36 and ODP849. Glacial periods (isotopic stages 2
and 4) are shown in light gray [Martinson et al., 1987]. See section 3.2 for information about the d18O
records. The vertical arrows in the upper panel represent the depth of the 14C ages that were used to
calculate the MD2138 age model (T. de Garidel-Thoron, manuscript in preparation, 2004). EEP, eastern
equatorial Pacific; WEP, western equatorial Pacific.
(IS1), 3 (IS3) and 5a (IS5a) (Figure 5). (231Pa/230Th)xs,0
is consistently higher at the western site.
4.4. The 230Th-Normalized Fluxes
[26] The total preserved rain rates are higher in the WEP
(1.4 g/cm2/ka), reflecting primarily higher fluxes of lithogenic material (Figure 6). Preserved carbonate rain rates in
MD2138 reach maximum values during IS1, IS3 and IS5a,
coincident to maxima in (231Pa/230Th)xs,0 (Figure 5). In the
EEP (ODP849), total preserved rain rates increase sharply
around 55 ka from 0.5 g/cm2/ka to 1.1 g/cm2/ka. This
variation is mainly due to a drastic change in preserved
carbonate rain rates, which may reflect either a change in
carbonate production in the overlying water or a change in
carbonate preservation on the seafloor. Our ODP849 preserved opal fluxes record (Figure 6) also suggests two
minima during IS2 and IS4, coincident with the two minima
in (231Pa/230Th)xs,0.
4.5. Syndepositional Sediment Redistribution (y
y)
[27] The sediment focusing factor (y) has been calculated
using equation (3), averaging the dry bulk density and
(230Thxs,0) over the four climatic stages delineated by the
d18O profiles (Table 4; Figure 7). Both cores show evidence
for sediment focusing (y > 1). y increases with time from
2 to 4 in MD2138 with slightly higher values during IS2
11 of 21
PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
PA4023
PA4023
Figure 6. 230Th-normalized fluxes (or preserved vertical rain rates) in cores MD2138 and ODP849.
Large symbols represent fluxes deduced from elemental ratios. The smaller high resolution (HR)
connected symbols for MD2138 represent estimates derived from 232Th (see text for explanations).
Glacial periods (isotopic stages 2 and 4) are shown in light gray [Martinson et al., 1987].
than during the Holocene. At ODP site 849, y is clearly
higher during IS2 and IS4.
5. Discussion
5.1. Authigenic U Accumulations in MD2138
[28] A surprising aspect of core MD2138 is its reducing
conditions below 60 cm depth that enabled the dissolution
and remobilization of MnO2 and the precipitation of authigenic U (Figure 4). We have estimated the depth (ZU) at
which the sediment had to become sufficiently reducing to
precipitate U, assuming that pore water U concentration
approaches zero at this depth [e.g., François et al., 1993]:
ZU ¼ DU ½U sw = SR r
auth
½U sed ;
ð6Þ
where DU is the pore water diffusion coefficient for U, 3
106 cm2/s [Klinkhammer and Palmer, 1991]; [U]sw is the
concentration of U in seawater, 1.3 1011 mol/cm3; SR is
the sedimentation rate, cm/s; r is dry bulk density, g/cm3;
and auth[U]sed is the authigenic U concentration in sediment,
mol/g.
[29] Applying equation (6) to the authigenic U profile in
MD2138 gives ZU 20– 35 cm. Using the DU value of
Mangini et al. [2001] (DU = 1.09 106 cm2/s) would give a
slightly shallower depth ZU 6– 12 cm. Authigenic U ceased
to accumulate above 70 cm depth, i.e., around 10 ka
(Figure 4), i.e., soon after the last deglaciation. Reducing
conditions in sediments result either from low oxygen
concentration in bottom water or from high flux of labile
organic matter to the sediment. MD2138 is located in a region
that is not characterized by oxygen-deficient intermediate or
deep waters at Present [Deuser, 1975]. Low oxygen concentrations in bottom waters of the WEP would imply a dramatic
slowdown of intermediate and deep water circulation in this
area. Studies have shown that Pacific intermediate and deep
Table 4. Sediment Focusing in Core MD2138 and ODP849a
Isotopic
Stage
Sedimentation
Rate, cm/kyr
Average
Carbonate, %
Average DBD,
g/cm3
Focusing
Factor, y
Holocene
IS2
IS3
IS4
IS5a
9.4
12.3
8.1
6.0
5.3
MD2138
62
52
64
59
57
0.52
0.44
0.55
0.50
0.48
3.6
4.1
3.7
3.0
2.0
Holocene
IS2
IS3
IS4
3.5
4.3
4.0
3.5
ODP849
72
77
71
61
0.62
0.69
0.62
0.52
2.1
3.0
1.8
2.7
a
Sedimentation rates were averaged on each isotopic stage. DBD, dry
bulk density.
12 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
Figure 7. Estimates of sediment focusing (y) as a result of
syndepositional sediment redistribution by bottom currents (see text for explanations). The horizontal dashed
line represents the limit between focusing (y > 1) and
winnowing (y < 1). Glacial periods (isotopic stages 2
and 4) are shown in light gray [Martinson et al., 1987].
water circulation has either remained similar to the Present
circulation or had an additional north Pacific component
during the LGM [Matsumoto and Lynch-Stieglitz, 1999;
Matsumoto et al., 2002]. It is therefore unlikely that bottom
waters overlying the site of MD2138 ever had very low
concentrations of oxygen resulting from a change in deep
water circulation. Hence high flux of labile organic matter
appears to be the most likely explanation for the reducing
conditions in MD2138. The timing of the change of the
reducing conditions in MD2138 suggests that this event is
associated with the last deglaciation. Core MD2138 was
collected on the continental slope of Manus Island. Sediment
accumulation rates are relatively high, particularly during the
last 24 ka (11 cm/ka; Table 4). 230Th accumulation indicates significant focusing of sediment (Figure 7). Low sea
level stand results in greater input of lithogenic material to the
deep sea [e.g., François and Bacon, 1991], and we do find a
significant (40 – 50%) decrease of the lithogenic flux during
the last deglaciation (Figure 6). Given MD2138 location, the
laterally transported sediments are likely to reach the coring
site by downslope transport from shallower depths, supplying enough labile organic matter to induce reducing conditions in core MD2138. Accumulation of authigenic U in
sediments as a result of sediment focusing has also been
observed in other oceanic regions [e.g., François et al.,
1993]. However, there are no clear changes in focusing factor
during deglaciation, which could readily explain the redox
transition (Figure 7). At this stage, we can only speculate that
the amount of organic matter associated with the sediment
transported downslope could have been higher before the last
deglaciation. Further studies will be needed to solve this
problem. In particular, it would be interesting to check
whether the redox conditions have changed at the last
deglaciation over the entire basin where MD2138 is located
or only locally.
5.2. (231Pa//230Th)xs in Surface Sediment
[30] In the Pacific Ocean, the residence time of deep
water (600 years) is much longer than the lateral diffusive
PA4023
mixing time (100 years [Anderson et al., 1990]), thus
allowing a full expression of boundary scavenging and a
greater sensitivity of (231Pa/230Th)xs,0 to particle flux [Yu et
al., 2001]. In order to further document the extent of
boundary scavenging in the Pacific ocean, we have compared (231Pa/230Th)xs in surface sediment with estimates of
export production. The latter was obtained from ocean color
measurements by satellite and a temperature-dependent
ecosystem model [Laws et al., 2000].
[31] In the EEP, we find a significant linear correlation
(R2 = 0.69) between (231Pa/230Th)xs measured in the surface
sediments and export production (Table 5; Figure 8). This
positive relationship suggests that (231Pa/230Th)xs,0 can be
used as a paleoproductivity proxy in the EEP.
[32] In the WEP, Laws et al. [2000] model predicts export
production lower than 12 gC m2 yr1 (Table 5). These
low values are inconsistent with the generally high
(231Pa/230Th)xs,0 ratios (>0.093) measured in the core top
of MD2138 and other cores from the surrounding area
(Table 5; Figure 8). Three explanations could account for
this discrepancy: (1) increased boundary scavenging resulting from higher influx of lithogenic material in the WEP
from surrounding continental masses [e.g., Milliman et al.,
1999], (2) underestimation of export productivity in the
WEP by ocean color measurements from satellites or
(3) control of the (231Pa/230Th)xs,0 by the chemical composition of the sinking particles.
[33] Since 232Th is exclusively associated with lithogenic material, it can be used as a proxy for lithogenic
fluxes. Thus hypothesis (1) can be tested by comparing
the surface (231Pa/230Th)xs to the 230Th-normalized 232Th
flux in the same cores. As expected, there is a general
eastward decreasing trend in 232Th flux in the WEP
(Table 6), reflecting a decrease in terrigenous flux away
from lands. A similar trend is also recognized in sediment
trap data [Kawahata et al., 2000]. This decrease is not
matched, however, by a decrease in (231Pa/230Th)xs,
which results in a total lack of correlation between
(231Pa/230Th)xs and the flux of 232Th, i.e., of lithogenic
material (Figure 9a). Increased boundary scavenging due
to high lithogenic flux fails to explain the high
( 231 Pa/ 230 Th) xs,0 measured in MD2138. One possible explanation for the lack of correlation between
(231Pa/230Th)xs and satellite-derived estimates of export
production in the WEP (Figure 8) is that the latter
underestimate productivity (hypothesis 2). The validity
of estimates of export productivity derived from ocean
color in the WEP has been challenged by a global-scale
comparison of organic carbon fluxes measured with deepsea moored sediment traps to export production derived
from Laws et al. [2000] satellite-based algorithm. Results
show that the latter might underestimate the settling flux
of organic carbon reaching the bathypelagic zone in some
areas which includes two data points from the WEP
[François et al., 2002]. Sediment trap measurements also
show that export biogenic fluxes are relatively high (8.5 –
47 g m2 yr1) in the WEP [Kawahata et al., 2000].
However, the sediment trap database is too sparse to
verify hypothesis (2). Thus if the WEP biogenic particle
fluxes are 10 gC m2 yr1 as predicted by Laws et al.
13 of 21
PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
PA4023
Table 5. Export Production (EP) Derived From Laws et al. [2000]
and (231Pa/230Th)xs,0 in the Surface Sediments From Sites in the
Equatorial Pacifica
Core
V19-29
P7*
V19-28
Y71-3-02
Y69-071P
KH-71-5-42-2
VNTR01-16PC
VNTR01-15GC*
VNTR01-19PC
VNTR01-13GC
KH-71-5-44-2
VNTR01-21GC
VNTR01-12GC
VNTR01-11GC
VNTR01-22GC
KLH 068*
VNTR01-10GC*
KLH 093*
MANOP M*
VNTR01-01PC
VNTR01-03GC
VNTR01-02PC
VNTR01-04GC
VNTR01-08PC
VNTR01-09GC
ODP849
VNTR01-06GC
VNTR01-07GC
VNTR01-05GC
TT154-10*
154-10*
KH-71-5-53-2
154-18*
154-8*
V19-55*
154-19
154-20
154-6
154-5
C57-58
154-4
MANOP C
B52-39
MC112
MC69
MC34
MC27
MC19
RC11-210
MANOP S
MC97
1858 358 bl
GIK10145-1
GIK10147-1
10175
KH-71-5-12-3
1858 163 bl
1858 232 bl
1858 195 bl
1858 254 bl
GIK10149-1
KH-71-5-15-2
GIK10140-1
Valdivia 10141
GIK10141-1
GIK10132-1
Longitude,
E; +W
Latitude,
S; +N
EP,
gC/m2/y
Eastern Equatorial Pacific
83.93
3.58
83.99
2.61
84.65
2.37
85.15
7.17
86.48
0.09
88.05
27.58
89.73
2.60
89.86
1.49
90.44
7.91
90.82
3.09
93.35
20.84
94.60
9.59
95.07
3.01
95.34
0.14
99.37
13.01
101.61
1.23
102.02
4.51
102.06
1.23
104.00
8.80
109.61
11.25
109.74
7.17
109.75
7.19
110.09
5.35
110.48
0.04
110.50
3.00
110.52
0.20
110.55
2.76
110.57
1.02
110.58
2.76
111.33
10.28
111.33
10.29
112.70
8.26
113.86
20.03
113.87
10.81
114.18
17.00
116.63
19.84
117.97
19.66
119.78
12.07
125.60
12.32
125.91
15.16
134.85
12.62
138.93
1.03
139.07
11.25
139.64
5.08
139.74
0.12
140.00
5.00
140.00
3.00
140.00
2.00
140.05
1.82
140.08
11.05
140.15
2.06
143.55
8.01
144.82
3.99
145.03
3.84
146.02
9.32
146.03
11.02
146.03
9.27
146.05
9.35
146.09
9.67
146.09
9.33
146.16
9.51
148.04
20.38
148.74
9.25
148.78
9.11
148.78
9.11
148.96
6.22
(EEP)
39.2
19.6
32.5
18.0
25.3
14.0
18.0
21.5
24.5
28.2
14.9
22.3
28.2
28.1
18.5
20.5
20.3
20.4
17.1
14.7
14.5
14.5
15.9
22.5
19.7
22.3
19.0
21.0
19.0
17.1
17.1
17.5
12.4
16.9
14.7
12.3
12.2
15.7
14.5
15.4
13.4
20.8
15.0
17.3
21.4
17.0
17.7
19.0
20.1
14.5
19.7
14.7
17.9
17.9
12.7
14.1
12.7
12.7
12.4
12.8
12.6
10.8
12.4
12.5
12.5
16.2
(231Pa/230Th)xs,0
0.177
0.250
0.200
0.121
0.193
0.054
0.106
0.167
0.136
0.143
0.025
0.150
0.119
0.137
0.154
0.167
0.128
0.181
0.159
0.056
0.089
0.073
0.093
0.113
0.109
0.116
0.092
0.116
0.094
0.157
0.160
0.080
0.199
0.110
0.148
0.076
0.057
0.077
0.070
0.034
0.037
0.068
0.033
0.048
0.070
0.060
0.050
0.060
0.071
0.027
0.060
0.034
0.027
0.037
0.017
0.091
0.027
0.034
0.023
0.031
0.035
0.034
0.037
0.028
0.029
0.032
PA4023
Table 5. (continued)
Core
V18-299
1858 21 bl
1858 151 bl
A47-16
KK1, core2
10127-2
Valdivia 10127
KK1, core1
V21-59
210KG
214KG
G993
V18-258
KH-79-1-5
MD2138
KH-79-4-6
KH-79-4-7
oj erdc bx88
KH-79-4-8
KH-79-4-9
KH-79-4-22
KH-79-4-10
V28-238
BC36
oj erdc bx125
KH-79-4-18
KH-79-4-19
KH-78-1-1036
KH-78-1-1038
Longitude,
E; +W
Latitude,
S; +N
EP,
gC/m2/y
(231Pa/230Th)xs,0
149.67
150.05
150.17
151.19
151.57
151.66
151.98
153.17
158.10
160.50
161.53
162.90
165.75
16.12
12.32
12.34
9.04
15.33
13.70
13.70
14.12
20.92
21.60
21.60
23.54
11.87
9.8
10.7
10.6
12.1
10.5
10.2
10.2
10.1
13.0
10.0
10.1
11.2
9.3
0.044
0.043
0.032
0.028
0.034
0.031
0.031
0.040
0.081
0.023
0.032
0.036
0.028
(WEP)
12.0
10.9
10.7
8.2
10.3
9.0
9.9
9.0
10.3
9.8
10.1
10.1
11.6
9.8
9.8
9.3
0.095
0.129
0.022
0.033
0.087
0.026
0.123
0.040
0.175
0.067
0.093
0.053
0.049
0.042
0.031
0.027
Western Equatorial Pacific
130.47
5.18
146.40
1.70
147.62
23.79
153.72
10.79
155.87
0.05
156.14
5.01
158.11
0.29
158.60
20.05
159.31
3.32
160.48
1.02
161.00
0.00
161.00
0.05
164.00
0.01
165.92
5.06
176.95
8.01
176.99
10.00
a
See Walter et al. [1999, and references therein]. Asterisks denote cores
located on/near the East Pacific Rise (EPR). Values are reported in Figure 8.
Parentheses denote activity.
[2000] model, then they cannot generate the high
(231Pa/230Th)xs measured in the surface sediment of the
area. The (231Pa/230Th)xs ratios could therefore be increased
by the preferential scavenging of 231Pa due to the chemical
composition of the sinking particles (hypothesis 3). Sediment trap experiments have shown that opal is produced
and exported in the WEP (about 20 – 30% of the total
particle flux [Kawahata et al., 2000]) but is not preserved
in the sediment. Chase et al. [2002, 2003a] have shown that
the particle reactivity of 231Pa increase with the opal content
and that 230Th has the opposite behavior. Opal fluxes could
therefore explain, at least partially, the high (231Pa/230Th)xs
ratios measured in the WEP sediments at Present. The
MnO2 recorded in the surface of MD2138 could also
explain the high (231Pa/230Th)xs ratios since MnO2 does
not fractionate 231Pa from 230Th. Therefore spatial variations in the opal and MnO2 fluxes could account for the
geographical (231Pa/230Th)xs variability recorded in the
surface of the sediments. However, this hypothesis cannot
be verified because of the lack of data. Alternatively, the
lowest opal/carbonate ratios in the WEP are found at the
equator [Kawahata et al., 2000] where the highest
(231Pa/230Th)xs ratios are measured (Figure 9b). This behavior is opposite to what is expected from Chase et al.
[2003a] results.
[34] To summarize, the high values of the (231Pa/230Th)xs
ratios measured in the surface sediments of the WEP could
either be due to the composition of the particles (opal- or
14 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
PA4023
Table 6. (231Pa/230Th)xs,0 and Fluxes of 232Th in the WEP as
Proxy for Lithogenic Input to the Sedimenta
230
Core
Figure 8. (231Pa/230Th)xs measured in the surface sediments of the eastern (EEP) and western (WEP) equatorial
Pacific versus export production estimated from ocean color
measurements by satellite and a temperature-dependent
ecosystem model [Laws et al., 2000]. Crosses indicate data
from cores collected in the vicinity of the East Pacific Rise
(EPR). They were not used in the regression because of
possible high MnO2 concentration, which would have
affected their (231Pa/230Th)xs. There is a positive linear
relationship between (231Pa/230Th)xs and export production
in the EEP (R2 = 0.69). Source data are in Table 5.
MnO2-rich or both) and/or to export fluxes higher than
those determined by satellite-based models. The spatial
variability of the (231Pa/230Th)xs could be due to variations
of the opal and/or MnO2 fluxes and/or to variations in the
export production. Testing these hypotheses would require
having much better constraints on the particle fluxes and
their chemical compositions.
5.3. Downcore Variations in (231Pa//230Th)xs,0
5.3.1. Local Versus Remotely Induced Variations of
the (231Pa//230Th)xs,0
[35] We first have to discuss whether the observed glacial
decrease in the (231Pa/230Th)xs,0 in the WEP and the EEP are
due to local or remote phenomenon. The generally stronger
glacial winds [e.g., Parkin and Shackleton, 1973; Sarnthein
et al., 1981] are likely to enhance the upwelling intensity on
the western margins of the continents which would increase
both productivity and eolian inputs in these areas. This
phenomenon could increase the scavenging of 231Pa at the
ocean margins therefore increasing the (231Pa/230Th)xs,0
ratios recorded in the underlying sediments. Consequently,
more 231Pa would be advected from the adjacent regions
which would be characterized by a decrease of the
(231Pa/230Th)xs,0 values during glacial times. Although the
downcore (231Pa/230Th)xs,0 database is very sparse for
the Pacific, the few available data [Lao et al., 1992] show
no significant glacial to Holocene (231Pa/230Th)xs,0 changes
at the western margins. The cores located at the eastern
margins (Californian and the equatorial South American)
show lower glacial (231Pa/230Th)xs,0 ratios. These results are
opposite to what is expected from enhanced glacial scavenging of 231Pa at the margins. Therefore the lower glacial
(231Pa/230Th)xs,0 measured in the WEP and the EEP appears
to be due to local phenomenon rather than to remotely
induced advection of 231Pa from the equatorial regions to
the margins.
KH-79-1-5
MD2138
KH-79-4-7
oj erdc bx88
KH-79-4-8
KH-79-4-9
KH-79-4-10
V28-238
BC36
oj erdc bx125
KH-79-4-18
KH-79-4-19
KH-78-1-1036
KH-78-1-1038
Longitude, Latitude,
E
S; +N (231Pa/230Th)xs,0
130.47
146.40
153.72
155.87
156.14
158.11
159.31
160.48
161.00
161.00
164.00
165.92
176.95
176.99
5.18
1.70
10.79
0.05
5.01
0.29
3.32
1.02
0.00
0.05
0.01
5.06
8.01
10.00
0.095
0.129
0.033
0.087
0.026
0.123
0.175
0.067
0.093
0.053
0.049
0.042
0.031
0.027
Th-Normalized
Th Fluxes,
mg/cm2/kyr
232
2.08
1.30
1.20
0.53
0.48
0.48
0.83
0.50
0.13
0.49
0.62
0.73
0.66
1.38
a
Values are reported in Figure 9.
5.3.2. Western Equatorial Pacific
[36] (231Pa/230Th)xs,0 is invariant on Ontong Java Plateau
(core BC36) over the entire Holocene, suggesting very little
changes in particle flux and composition at this site during
the last 10 ka. The situation is different at the westernmost
MD2138 site where we find a sharp mid-Holocene
(231Pa/230Th)xs,0 maximum (Figure 5). This longer record
Figure 9. (a) (231Pa/230Th)xs measured in the surface
sediments of the western equatorial Pacific versus 230Thnormalized 232Th fluxes used as a proxy for the fluxes of
lithogenic material. (b) (231Pa/230Th)xs measured in the
surface sediments (Table 6) and opal/carbonate recorded in
sediment traps [Kawahata et al., 2000] in the western
equatorial Pacific versus latitude. There is a decrease of the
(231Pa/230Th)xs values from the equator to the north.
15 of 21
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Figure 10. (231Pa/230Th)xs,0 versus (a) 230Th-normalized
carbonate fluxes and (b) 230Th-normalized lithogenic fluxes
measured downcore in MD2138.
also documents clear minima during the colder IS2 and IS4,
and higher values during IS3 and IS5a.
[37] The lack of constraints on the paleoceanography of
the WEP and the complex geological and oceanographic
settings of the region render the finding of an unique
explanation that could account for the (231Pa/230Th)xs,0
variations difficult. In the following paragraphs, we discuss
various hypotheses that could explain the lower glacial
(231Pa/230Th)xs,0 ratios measured in MD2138.
[38] As proposed in section 5.1., the variations of the
(231Pa/230Th)xs,0 ratio in MD2138 could be due to a change
in the chemical composition of the settling particles. The
high sediment focusing factors observed in MD2138
(Table 4) could be explained by downslope transport given
that the core is located at the foot of a continental slope.
Particles with relatively high MnO2 concentrations could
have been resuspended from shallower depths during the
downslope transport. MnO2-rich particles scavenge dissolved 230Th and 231Pa with the same efficiency. Consequently, the (231Pa/230Th)xs,0 ratios in MD2138 would be
increased if the resuspended particles have spent a substantial amount of time in the water column before their final
burial at the core site. If sediment focusing is interpreted in
terms of downslope transport, when sediment focusing
increase, higher fluxes of MnO2 are expected and accordingly higher (231Pa/230Th)xs,0 ratios. However, the continuous increase of sediment focusing from IS5 to IS3 (Figure 7)
is not matched by an increase of the (231Pa/230Th)xs,0
ratios (Figure 5). Similarly, from IS2 to the Holocene,
sediment focusing decreased and (231Pa/230Th)xs,0 ratios
increased while the opposite behavior is expected. The
downslope transport of MnO2-rich particles could therefore explain that (231Pa/230Th)xs,0 values in MD2138 are
generally higher than 0.093 but this phenomenon could
not account for the temporal variations of the
(231Pa/230Th)xs,0 ratios.
[39] The (231Pa/230Th)xs,0 maxima found in the Holocene,
IS3 and IS5a coincide with maxima in 230Th-normalized
PA4023
carbonate flux (Figure 6), suggesting a productivity control.
The relatively shallow depth at which this core was taken
(1900 m) is well above the calcite saturation horizon in the
water column of this region (3000 m) and likely to be
outside the range of depth variability of the saturation
horizon over the time considered. Therefore variability in
sedimentary calcite fluxes in this core should mainly reflect
variations in carbonate production unaffected by variations
in carbonate preservation. In addition, there is a weak
correlation between downcore variations in (231Pa/230Th)xs,0
and carbonate fluxes (R2 = 0.24), while none exist with
lithogenic fluxes (Figure 10). The (231Pa/230Th)xs,0 record of
MD2138 thus suggests lower carbonate export production
and particle flux in the WEP during glacial periods. Chase
et al. [2002, 2003a] have recently argued that the
(231Pa/230Th)xs,0 variability is mainly controlled by variations in the opal/carbonate ratio of the sinking particles. We
cannot rule out the possibility that, in addition to lower
export production by carbonate-producing phytoplankton,
(231Pa/230Th)xs,0 may also have been partly lowered by a
decrease of the opal/carbonate ratio of the sinking particles
which then would correspond to a drop in diatom production. There are only negligible amounts of preserved opal
in MD2138. However, at Present, the biogenic opal
produced in the WEP (20 – 30% of the vertical particle
flux [Kawahata et al., 2000]) is not preserved in the
sediment. If the same behavior holds for the past, the
high (231Pa/230Th)xs,0 ratios could also be explained by
the preferential scavenging of 231Pa relative to 230Th
induced by opal [Chase et al., 2002, 2003a]. Further
studies are needed to constrain the variations of the
chemical composition and intensity of the biogenic fluxes.
However, our study tends to show a decrease of the
exported carbonate, and possibly opal, fluxes during IS2
and IS4 in the WEP.
5.3.3. Eastern Equatorial Pacific
[40] (231Pa/230Th)xs,0 values are generally lower at ODP
site 849 compared to the western site (MD2138). However,
the variability in the ratio is very similar in both cores, with
lower values during the colder periods (IS2 and IS4) and
higher values during the warmer periods (Holocene, IS3 and
IS5a). The main difference is a broader Holocene maximum
in the eastern Pacific core (Figure 5). The slight delay found
during IS4 could result from uncertainties in the
core chronology. In contrast to core MD2138, the
(231Pa/230Th)xs,0 profile in ODP849 does not fully mimic
the variations of the 230Th-normalized carbonate flux
(Figure 6). The sharp increase of (231Pa/230Th)xs,0 between
60 and 50 ka is associated to an increase of the 230Thnormalized carbonate flux. However, this is not true for the
(231Pa/230Th)xs,0 increase at the last deglaciation. ODP849
was collected at greater depth (3800 m) than MD2138, well
below the saturation horizon. Therefore variations in the
preserved carbonate rain rate reconstructed by 230Th normalization not only reflect changes in carbonate production,
but also changes in carbonate preservation resulting from
changes in the depth of the lysocline [e.g., Farrell and Prell,
1989]. In further contrast to the western site, some opal is
preserved in sediments deposited at the ODP849 site
(Figure 3). The variability in preserved opal rain rates
16 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
Figure 11. (231Pa/230Th)xs,0, 230Th-normalized preserved
opal fluxes, and carbonate production rate in ODP849.
Carbonate production rates were obtained by combining
230
Th-normalized preserved carbonate fluxes and quantitative estimates of carbonate dissolution from a G. menardii
fragmentation transfer function [Loubere et al., 2004].
measured in this core is very similar to the (231Pa/230Th)xs,0
record (Figure 11). This observation and the broad correlation that we find between satellite-derived export production and (231Pa/230Th)xs in the surface sediment of the EEP
(Figure 8) suggest that the low glacial (231Pa/230Th)xs,0 is
due to lower export productivity. Lower carbonate production rates during IS2 were also obtained from the same
core (Figure 11) by Loubere et al. [2004] after combining
230
Th-normalized preserved carbonate fluxes and quantitative estimates of carbonate dissolution from a G. menardii
fragmentation transfer function [Mekik et al., 2002]. Both
approaches point to lower glacial productivity, for
both carbonate and opal producing plankton, at ODP Site
849.
5.4. Comparison With Previous Studies of
Paleoproductivity in the Equatorial Pacific
[41] (231Pa/230Th)xs,0 (Figure 5), calcite production rates
[Loubere et al., 2004], and 230Th-normalized fluxes of
preserved opal (Figure 11) consistently suggest lower glacial productivity in the western and eastern equatorial
Pacific. Lower glacial productivity in the EEP is consistent
with recent paleoproductivity reconstructions in this region
using a new transfer function based on benthic foraminifera
assemblages [Loubere, 1999, 2000, 2001, 2003], but contrary to conclusions derived from mass accumulation rates
of biogenic material [Lyle et al., 1988; Sarnthein et al.,
1988; Paytan et al., 1996]. It is becoming increasingly
evident that mass accumulation rates (MAR) of biogenic
material on the seafloor can be significantly affected by
sediment redistribution by bottom currents [Marcantonio et
al., 2001; François et al., 2004; Loubere et al., 2004]. In
particular, the study by Marcantonio et al. [2001] showed
that most of the variability in barite MAR in the central
equatorial Pacific is eliminated by normalizing the barite
flux to 230Thxs,0. Preserved calcite MAR in ODP849 is
significantly higher at the LGM than during the Holocene.
However, Loubere et al. [2004] study showed that the
difference vanishes when using 230 Th xs,0 -normalized
PA4023
calcite flux. The MAR values calculated in ODP849
(1.7 – 2.4 g/cm2 /ka) are 2 – 3.5 times higher than the
230
Th xs,0 -normalized total flux (0.5 – 1.2 g/cm 2 /ka)
(Figure 12). In particular, during IS2, 230Thxs,0-normalized
total flux decreases by 15% whereas MAR exhibits a 25%
increase. This latter appears to be mostly due to the glacial
increase of sediment focusing (Figure 7). These studies
suggest that variations in MAR in the equatorial Pacific
mainly reflect variability in sediment focusing rather than
changes in particle flux from the overlying surface water. Our
(231Pa/230Th)xs,0 profiles together with the high focusing
values we calculated in ODP849 support this interpretation
and point out the necessity of using 230Thxs,0-normalization
rather than MAR to reconstruct export fluxes.
5.5. Implications From Lower Glacial Productivity in
the WEP and the EEP
[42] Lower glacial productivity shown in our study could
suggest that the El Nino climatic mode would be more
prominent during glacial periods. This interpretation is
consistent with smaller SST gradients and lower equatorial
upwelling rates reported by Koutavas et al. [2002] and
modeling studies of Clement et al. [1999]. However, irrespective of the intensity of the Trade Winds and the
equatorial upwelling rate, lower productivity could also
result from lower nutrient concentrations in the EUC which
supply nutrient to the equatorial divergence and the SEC
[e.g., Loubere, 1999; Spero and Lea, 2002]. Various
hypotheses could account for lower nutrient concentrations
in the EUC. For instance, upwelling of Fe-rich waters from
the EUC has been proposed as a major source of Fe to
the EEP [Gordon et al., 1997]. Fe in the EUC originates
from the interaction between the New Guinea coastal
undercurrent (NGCU) and the continental shelf of Papua
New Guinea [Mackey et al., 2002] where great loads of
continental inputs occur (860 t/yr [Milliman et al., 1999]).
During the last glacial maximum the seawater level has
dropped by 100– 120 m [Lambeck and Chappell, 2001, and
references therein]. Consequently, the interaction between
Figure 12. 230Th-normalized total flux (or preserved
vertical rain rate) and marine accumulation rate (MAR)
averaged over each isotopic stage in ODP849. Glacial
periods (isotopic stages 2 and 4) are shown in light gray
[Martinson et al., 1987].
17 of 21
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PICHAT ET AL.: LOWER GLACIAL PRODUCTION IN THE PACIFIC
the NGCU and the Papua New Guinea continental shelf
could have been reduced potentially bringing less Fe into
the EUC. Recently, Loubere [1999], Matsumoto et al.
[2002] and Brzezinski et al. [2002] have argued for lower
nitrate and higher silicate concentrations in the glacial EUC.
They proposed a Fe-induced increase in the nitrate/silicate
uptake ratio in the surface water of the Southern Ocean from
where the deeper part of the EUC originates [Toggweiler et
al., 1991; Rodgers et al., 2003]. They postulate that such an
increase could have resulted in a higher supply of silicate to
the EEP via the EUC thereby increasing diatom production
and the organic carbon/carbonate rain ratio. Consequently,
the surface water alkalinity would have increased thus
contributing to the glacial drawdown of atmospheric CO2
(silicic acid leakage hypothesis). Sea surface cooling in the
tropics would have also provided an abiotic contribution to
this drawdown. The (231Pa/230Th)xs,0 and 230Th-normalized
fluxes of preserved opal from ODP849 suggest lower
diatom productivity during glacial periods, which seems
to challenge the silicic acid leakage hypothesis. However,
deriving a conclusion from a single core may be premature
given the large regional variability of the productivity
response to glacial climate that characterize the EEP
[Loubere, 2000, 2003]. At ODP Site 846 (3.1S, 90.8W),
located west of ODP849 under the SEC, there are preliminary evidence for higher opal preserved rain rate during the
last glacial period, higher rain ratio and lower overall
productivity (P. Loubere et al., manuscript in preparation,
2004). Fully evaluating the validity of the silicic acid
leakage hypothesis and verifying whether El Nino is an
adequate description of the oceanographic setting of the
glacial equatorial Pacific will therefore require detailed
synoptic reconstructions of primary production and SST
over the entire equatorial Pacific region.
6. Conclusions
This study shows similar variations in ( 231Pa/
Th) xs,0 in two sediment cores from the western
(MD2138) and eastern (ODP849) equatorial Pacific with
systematically lower values during isotopic stages 2 and 4,
i.e., glacial periods. Given the lack of data and constraints
on the paleoceanography of the western equatorial Pacific,
the conclusions drawn from our study of core MD2138
are still ambiguous. The generally high (231Pa/230Th)xs,0
[43 ]
230
PA4023
ratios measured in MD2138 could be due to opal fluxes
that are not preserved in the sediment and/or MnO2 fluxes
of downslope transported particles. Alternatively, although
this hypothesis is questionable with Present data, the
export production could have been high enough to explain
the generally high (231Pa/230Th)xs,0. Variations in downslope transport of MnO2 particles cannot explain the
variations in the (231Pa/230Th)xs,0 profile. The study of
230
Th-normalized carbonate flux variations show a decrease
in export biogenic carbonate that could explain the glacial
decrease of (231Pa/230Th)xs,0. Lower opal/carbonate ratios,
i.e., a decrease in the exported opal flux, could also account
for the lower (231Pa/230Th)xs,0 glacial values. Although in
need of confirmation, our study tends to show a decrease of
export production (carbonate and/or opal) during glacial
periods in the western equatorial Pacific.
[44] Combined with profiles of elemental composition
and 230Thxs-normalized fluxes, the (231Pa/230Th)xs,0 variations in the eastern equatorial Pacific can be interpreted
as reflecting lower export production during glacial periods. Our conclusions for the EEP are in agreement with
recently developed proxies that are insensitive to dissolution or sediment redistribution processes [Loubere et al.,
2004].
[45] There is significant sediment focusing at the two
study sites. Particularly, there is higher focusing during
glacial periods in the eastern equatorial Pacific. Our
results, together with conclusions from previous studies
[Marcantonio et al., 2001; Loubere et al., 2004] highlight
the necessity of using 230Th-normalization instead of mass
accumulation rates (MAR) to reconstruct past changes in
export flux, in particular biogenic paleofluxes.
[46] Acknowledgments. We thank Lary Ball and the WHOI ICP
Facility for use of their Finnigan MAT ELEMENT SF-ICP-MS and their
Jobin Yvon JY38VHR ICP-OES. We thank Steve Manganini for
performing analyses of oxides. Alan Fleer is thanked for his help in
the laboratory. We thank Luc Beaufort for his help while sampling
IMAGES core MD97-2138 at the CEREGE (France) core repository and
Dan McCorkle for his help at the WHOI (USA) core repository. Bob
Anderson, Marcus Christl, Gideon Henderson and Augusto Mangini are
thanked for their very constructive reviews. SP funding for this research
was provided by grants from the French Minister of Research and a
EURODOC grant of the Région Rhône-Alpes (SAFIR-980065327). SP
also gratefully acknowledges the financial support of the WHOI Geology
and Geophysics Dept. This work was also supported by a CNRS-NSF
grant (SP and KWWS). The contribution of JFM to this study was
supported in part by the US NSF and by WHOI OCCI and Mellon
awards. This is WHOI contribution 11054.
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S. Brown Leger and R. François, Department
of Marine Chemistry and Geochemistry, Woods
Hole Oceanographic Institution, Woods Hole,
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J. F. McManus and Kenneth W. W. Sims,
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S. Pichat, Department of Earth Sciences,
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3PR, UK. (sylvainp@earth.ox.ac.uk)