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8.16 Tracers of Past Ocean CirculationJ Lynch-Stieglitz, Georgia Institute of Technology, Atlanta, GA, USATM Marchitto, University of Colorado, Boulder, CO, USA

ã 2014 Elsevier Ltd. All rights reserved.

This article is a revision of the previous edition article by J. Lynch-Stieglitz, volume 6, pp. 433–451, © 2003, Elsevier Ltd.

8.16.1 Introduction 4358.16.2 Nutrient Water Mass Tracers 4368.16.2.1 Carbon Isotopes 4368.16.2.1.1 Controls on d13C of oceanic carbon 4368.16.2.1.2 Carbon isotope ratios in benthic foraminifera 4378.16.2.2 Cadmium in Benthic Foraminifera 4388.16.2.2.1 Cd in seawater 4388.16.2.2.2 Cd/Ca in foraminifera 4398.16.2.3 Ba/Ca 4408.16.2.4 Zn/Ca 4408.16.2.5 Silicon Isotopes 4408.16.3 Conservative Water Mass Tracers 4408.16.3.1 Mg/Ca in Benthic Foraminifera 4408.16.3.2 Pore Water Chemistry 4408.16.3.3 Oxygen Isotopes in Benthic Foraminifera 4418.16.3.4 Clumped Isotopes 4418.16.3.5 Carbon Isotope Air–Sea Exchange Signature 4418.16.4 Neodymium Isotope Ratios 4428.16.5 Circulation Rate Tracers 4428.16.5.1 Radiocarbon 4428.16.5.2 231Pa/230Th Ratio 4438.16.5.3 Geostrophic Shear Estimates from d18O in Benthic Foraminifera 4438.16.6 Nongeochemical Tracers of Past Ocean Circulation 4438.16.7 Ocean Circulation during the LGM 4448.16.8 Conclusions 448References 448

8.16.1 Introduction

Information about how the ocean circulated during the past is

useful in understanding changes in ocean and atmospheric

chemistry, changes in the fluxes of heat and freshwater be-

tween the ocean and the atmosphere, and changes in global

wind patterns. The circulation of surface waters in the ocean

leaves an imprint on sea surface temperature and is also inex-

tricably linked to the patterns of oceanic productivity. Much

valuable information about past ocean circulation has been

inferred from reconstructions of surface ocean temperature and

productivity, which are covered in separate chapters. Here, the

focus is on geochemical tracers that are used to infer the flow

patterns and mixing of subsurface water masses.

Several decades ago, it was realized that the chemistry of the

shells of benthic foraminifera (carbon isotope and Cd/Ca

ratios) carries an imprint of the nutrient content of deep water

masses (Boyle, 1981; Broecker, 1982; Shackleton, 1977).

This led rapidly to the recognition that water masses in the

Atlantic Ocean were arrayed differently during the Last Glacial

Maximum (LGM) than they are today and the hypothesis that

the glacial arrangement reflected a diminished contribution of

atise on Geochemistry 2nd Edition http://dx.doi.org/10.1016/B978-0-08-095975

low-nutrient North Atlantic Deep Water (NADW; Boyle and

Keigwin, 1982; Curry and Lohmann, 1982). More detailed

spatial reconstructions indicated a shallow nutrient-depleted

water mass overlying a more nutrient-rich water mass in the

glacial Atlantic. These findings spurred advances not only in

geochemistry but also in oceanography and climatology, as

workers in these fields attempted to simulate the inferred glacial

circulation patterns and assess the vulnerability of the modern

ocean circulation to changes observed for the last ice age.

While the nutrient distributions in the glacial Atlantic

Ocean were consistent with a diminished flow of NADW,

they also could have reflected an increase in inflow from the

South Atlantic and/or a shallower yet undiminished deep

water mass. Clearly, tracers capable of giving information on

deep water flow rate, rather than nutrient content alone, were

needed. Differences between surface water (measured on

planktonic foraminifera) and deepwater (measured on coexist-

ing benthic foraminifera) radiocarbon concentrations pro-

vided the first rate constraint (Broecker et al., 1988;

Shackleton et al., 1988). Reduced amounts of Pa relative to

the more particle-reactive Th in the glacial Atlantic suggested

that deep water was exported from the Atlantic during glacial

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http://dx.doi.org/10.1016/B978-0-08-095975-7.00616-1

436 Tracers of Past Ocean Circulation

times (Yu et al., 1996). Density gradients in upper waters

reflect the vertical shear in oceanic transports, especially at

intermediate depths where these gradients are largest (Lynch-

Stieglitz et al., 1999b).

Many of these tracers of paleo-ocean flow have been ap-

plied to all of the ocean basins and have been extended in time

throughout the Neogene. Despite this progress, a consistent

picture of the circulation of the ocean during even the last ice

age is yet to emerge. While circulation tracers suggest a rearran-

gement of water masses in the Atlantic, there is still consider-

able disagreement about the water masses and circulation in

the rest of the world ocean. Some of this uncertainty arises

from still insufficient data coverage but some is the result of

conflicting information from the various deep circulation

tracers. In this chapter, the methods used to reconstruct

past ocean circulation, which will illuminate the source of

some of this conflicting information, are examined in more

detail.

8.16.2 Nutrient Water Mass Tracers

8.16.2.1 Carbon Isotopes

Measuring the carbon isotope ratio in calcite tests of bottom-

dwelling (benthic) foraminifera is perhaps the most

widespread method for reconstructing the distribution and

properties of deep-sea water masses. Carbon isotope ratios in

foraminifera reflect that ratio in seawater. In turn, the carbon

isotope ratio in the deep sea is primarily controlled by the

regeneration of 13C-poor organic material and has a distribu-

tion much like a nutrient such as phosphate in the modern

ocean. While a nutrient tracer reflects the increased contribu-

tion from the regeneration of organic material as the water

mass ages, it also carries a signature of the initial nutrient

NADW

1.5

1.0

0.5

0.0

−0.51.0

PO4 (μ

d 13Ca

d 13Cas =0

δ13C

(‰ P

DB

)

Air–

sea

exch

angeBiology

1.5

Figure 1 Deepwater (>2000 m) d13C and PO4 data from the modern oceanfrom the mean ocean Redfield slope for d13C and PO4 (Broecker and Maier-Rslope expected for biologic processes, deep waters which form in the North Aexchange. Reproduced from Lynch-Stieglitz J, Stocker TF, Broecker WS, andcomposition of oceanic carbon: Observations and modeling. Global Biogeoch

content of the source regions. In today’s Atlantic, the contrast

between high-nutrient Southern Ocean source water and

low-nutrient NADW is strong, and the residence time of deep

water in the Atlantic Ocean is short, allowing little additional

accumulation of nutrients from sinking organic material. In

this case, the nutrient tracers can be used to document the

relative mixtures of northern and southern water masses.

8.16.2.1.1 Controls on d13C of oceanic carbonDuring photosynthesis, organisms preferentially take up the

lighter isotope of carbon (12C), increasing surface ocean d13C.When the 13C-depleted organic matter is decomposed, the

accompanying CO2 release decreases the d13C of dissolved

inorganic carbon at depth (Craig, 1970; Deuser and Hunt,

1969; Kroopnick, 1985). If this process were solely responsible

for the distribution of d13C in the ocean, d13C of oceanic

carbon would decrease by approximately 1.1% for every

1 mmol kg�1 increase in oceanic PO4 (Broecker and Maier-

Reimer, 1992). Reflecting this biological cycling, as a deep

water mass with a homogeneous source ages (such as in the

deep Indian and Pacific oceans), the d13C and PO4 increase in

this biologically expected ratio (Figure 1). However, the dis-

crimination against 13C during photosynthesis increases from

about 19% in the warm surface ocean to about 30% in the

Antarctic, which means that the relative changes in PO4 and

d13C associated with the uptake and regeneration of organic

material can vary regionally.

Carbon isotope fractionation during air–sea exchange is

also an important factor in determining the isotopic composi-

tion of seawater. If the CO2 in the atmosphere were in isotopic

equilibrium with the dissolved inorganic carbon in the ocean,

the dissolved inorganic carbon would be enriched in 13C rela-

tive to atmospheric CO2 by about 8% at 20 �C (Inoue and

Sugimura, 1985). This equilibrium fractionation depends on

WSBW

CDW

2.5mol kg−1)

d 13Cas =−0.5

s =0.5

2.0 3.0

. The lines indicate a ‘Redfield slope,’ and d13Cas indicates the deviationeimer, 1992). While the deep Indian and Pacific values follow thetlantic and Southern Ocean have lower and higher d13C due to air–seaFairbanks RG (1995) The influence of air–sea exchange on the isotopicemical Cycles 9: 653–665.

Figure 1

−0.40−0.40 1.20

Estimated seawater d13C (‰ PDB)

d13C

rec

onst

ruct

ed fr

om fo

ram

inife

ra (‰

PD

B)

0.800.400.00 1.60

1.20

0.80

0.40

0.00

1.60

Figure 2 Carbon isotopic composition of core-top benthic foraminiferain the genera Cibicidoides and Planulina versus estimated carbonisotopic composition of dissolved inorganic carbon in the water columnabove core location. Reproduced from Duplessy J-C, Shackleton NJ,Matthews RK, et al. (1984) 13C record of benthic foraminifera in the lastinterglacial ocean: Implications for the carbon cycle and the global deepwater circulation. Quaternary Research 21: 225–243.

Tracers of Past Ocean Circulation 437

the temperature of equilibration, with the oceanic carbon

more enriched relative to the atmospheric value by about

0.1% per degree of cooling (Mook et al., 1974). If the surface

ocean were everywhere in isotopic equilibrium with atmo-

spheric CO2, the 30 �C range in ocean temperatures would

cause a 3% range in oceanic d13C, with higher values where

surface temperatures are cold and lower values in warm surface

waters. This range is similar to the magnitude of d13C change

induced by biological processes. Isotopic equilibrium is rarely

reached, however, because the timescale for carbon isotope

equilibration between ocean and atmosphere (about 10 years

for a 50 m deep mixed layer) is about ten times longer

than chemical equilibration between atmospheric CO2 and

dissolved inorganic carbon (Broecker and Peng, 1974; Tans,

1980). This isotopic equilibration time is significantly longer

than the time that surface waters stay in one place in the ocean.

The exchange of atmospheric CO2 with surface ocean aque-

ous CO2 has the potential to leave oceanic carbon depleted in13C in areas of CO2 invasion (e.g., high northern latitudes,

low-pCO2 surface waters) and enriched in 13C in areas of

CO2 outgassing (e.g., equatorial upwelling regions, high-

pCO2 surface waters; Lynch-Stieglitz et al., 1995). This is be-

cause CO2 gas that is dissolved in the ocean and exchanges

between the ocean and atmosphere is isotopically lighter than

the much larger pool of dissolved inorganic carbon that con-

sists mainly of bicarbonate and carbonate ions. This effect is

observed in the ocean because the timescale for carbon equi-

librium between ocean and atmosphere (which drives the air–

sea exchange of CO2) is shorter than the timescale for isotopic

equilibrium, as noted above.

In general, surface ocean d13C lies somewhere between the

value predicted if biological cycling alone controlled the dis-

tribution of d13C in the ocean and the value that would be at

equilibrium with the atmosphere (Broecker and Maier-Reimer,

1992; Broecker and Peng, 1982; Gruber et al., 1999). The

departure from the biologically expected value can be

expressed as an air–sea exchange signature d13Cas¼d13C�(2.7�1.1�PO4) (Broecker and Maier-Reimer, 1992; Lynch-

Stieglitz et al., 1995). By definition, water with d13Cas of

0% has the same air–sea exchange signature as mean ocean

deep water. A positive value of d13Cas means that the water has

a higher d13Cas than mean ocean deep water (more of an

influence of air–sea exchange at cold temperatures/less at

warm temperatures), whereas a negative d13Cas value implies

less d13C enrichment due to air–sea exchange than for mean

ocean deep water.

d13Cas in surface waters is highest in the Antarctic and the

North Pacific due to exchange with the atmosphere at low

temperatures. The lowest values of d13Cas are found in the

centers of the subtropical gyres. Surface waters circulate in

these anticyclonic gyres long enough to allow the oceanic

DIC time to equilibrate isotopically with the atmosphere at

high temperatures, despite the relatively low rates of gas ex-

change. The situation in the North Atlantic seems slightly more

complex. Air–sea exchange at cold temperatures will tend to

raise the d13Cas of the northward moving upper waters but

does not have time to overcome the effects of CO2 invasion

and the remnant low d13Cas of the northward flowing surface

water. Deep waters leaving the surface in the North Atlantic

and the Antarctic reflect the low and high d13Cas, respectively.

8.16.2.1.2 Carbon isotope ratios in benthic foraminiferaBelanger et al. (1981) demonstrated that the d13C in the ben-

thic foraminifera Planulina wuellerstorfi decreased from highest

values in the Norwegian–Greenland Sea to lower values in the

South Atlantic and lowest values on the East Pacific Rise. They

argued that P. wuellerstorfi is a reliable recorder of bottom water

conditions and is not overly influenced by the incorporation of

metabolic carbon. (While wuellerstorfi is variously assigned to

the genera Planulina, Cibicides, Cibicidoides, and Fontbotia by

various workers, there is no disagreement on the assignment

of the species.) They also found that d13C in the genera

Oridorsalis and Pyrgo did not follow the expected deepwater

progression in d13C and argued that the shell chemistry

reflected the lower d13C of their pore water habitat. The same

year, Graham et al. (1981) showed, with a larger set of core-top

data, that d13C from the tests of P. wuellerstorfi and Cibicidoides

kullenbergi also follow the expected relationship with seawater

chemistry. Graham et al. (1981) also found that these species

show d13C values that are closest to what would be expected for

precipitation in seawater. While they found that d13C values in

P. wuellerstorfi and C. kullenbergi reflect the d13C of the ambient

water, these shells are a full 1% lower than expected for calcite

formed in equilibrium with this water (Woodruff et al., 1980).

However, carbon isotope equilibrium between calcite and dis-

solved carbon is not well known at the temperatures of the

deep ocean (Grossman, 1984b; Romanek et al., 1992). These

studies paved the way for the widespread reconstruction of

deepwater d13C which followed soon afterward (Boyle and

Keigwin, 1982; Curry and Lohmann, 1982; Curry et al., 1988;

Duplessy et al., 1984, 1988).

Shown in Figure 2 is a comparison of Planulina and Cibici-

doides d13C with estimated water column d13C from Duplessy

et al. (1984). At the time of this study, there was little

Figure 2


high-quality water column d13C data with which to compare

the core-top values. In subsequent years, however, more exten-

sive and improved datasets have confirmed that, on average,

Planulina and Cibicidoides d13C reflect water column d13C with

little or no offset. Some of the scatter in Figure 2 may be

attributed to errors in estimated oceanic d13C or core tops

which are not Holocene in age. However, the more recent

datasets also show that there can be significant offsets between

the foraminifera and water column values at some locations,

even when living foraminifera are compared to water column

d13C collected at the same location.

Perhaps it is not unexpected that benthic foraminifera do

not always cleanly record the isotopic composition of bottom

water. The fact that all of the benthic foraminifera have d13Cthat is significantly lighter than what would be expected

for equilibrium with bottom water provides the first clue that

additional processes may be at work. As postulated by Belanger

et al. (1981), some foraminifera may live in the pore waters of

the sediment, where d13C of dissolved inorganic carbon is

lower than in the overlying water column. Even so, the carbon

isotopic composition of the calcite skeleton of most calcitic

organisms does not reflect the composition of the ambient

waters (Wefer and Berger, 1991) but is dominated by physio-

logical or ‘vital’ effects which usually involve the incorporation

of isotopically light metabolic carbon or kinetic fractionation

during the transport of carbon to the calcification site or during

calcification itself. The isotopic ratios of planktonic foraminif-

era are also dominated by these vital effects and are, in general,

not in equilibrium with seawater (e.g., Spero and Lea, 1996;

Spero et al., 1997).

A more detailed look at how benthic foraminifera record

bottom water d13C has been made through careful analysis of

stained (live or recently living) benthic foraminifera collected

from surface sediment, combined with measurements of

both water column and pore water d13C (Grossman, 1984a,b;

Mackensen and Douglas, 1989; Mackensen et al., 1993;

McCorkle et al., 1990, 1997; Rathburn et al., 1996; Tachikawa

and Elderfield, 2002). These studies confirm earlier notions

that different species show different microhabitat preferences,

with some species (including most Cibicidoides and Planulina)

living on or near the sediment surface and others (including

Uvigerina) living within the sediment. The deeper-dwelling

species show depleted d13C values relative to bottom waters

that, in general, reflect the magnitude of the pore water de-

pletions. These studies suggest that this microhabitat effect is

the dominant control on benthic foraminiferal d13C and that

variations in the vital effects between species are small. While

the number of Cibicidoides analyzed has been small, the Cibici-

doides species tend to have the smallest isotopic depletion

relative to bottom waters.

However, the isotopic depletion even for Cibicidoides can be

significant, especially in areas where productivity, and thus the

supply of organic material to the sediment, is high. Mackensen

et al. (1993) showed that living and dead specimens of

P. wuellerstorfi showed substantial depletions relative to bot-

tom waters, particularly beneath high-productivity fronts in

the Southern Ocean. While P. wuellerstorfi has been reported

to live above the sediment surface (Lutze and Thiel, 1989),

Mackensen et al. (1993) postulate that the foraminifera are

calcifying within the organic fluff layer that can be found in

high-productivity environments, with their low d13C value

reflecting either the low-d13C environment within the fluff

layer or high growth rates of the foraminifera within these

environments. Subsequent studies, however, have shown that

the relationship between d13C and overlying productivity can

be complex (Corliss et al., 2006; Eberwein and Mackensen,

2006; Zarriess and Mackensen, 2011).

These relationships have been generally confirmed with

downcore d13C studies. These studies show that the d13Cfrom Uvigerina (infaunal) is always more negative than for

Cibicidoides and that this offset is variable in space and time,

consistent with the microhabitat model. While many early

isotopic studies were performed using Uvigerina because

it was thought to best reflect oxygen isotope equilibrium,

most deepwater d13C reconstructions now use Cibicidoides

or Planulina species. It is broadly acknowledged that these

records, particularly in high-productivity areas, might be influ-

enced by a microhabitat effect. However, by avoiding high-

productivity areas and using supplemental information about

changes in bottom water characteristics and overlying produc-

tivity, this method has continued to yield regionally coherent

and sensible patterns of deepwater d13C.

8.16.2.2 Cadmium in Benthic Foraminifera

The use of Cd/Ca measurements in the tests of benthic forami-

nifera for deepwater nutrient reconstructions was developed in

parallel with the carbon isotope method. Dissolved Cd con-

centrations in seawater follow a nutrient-like distribution,

while Ca concentrations mainly reflect variations in salinity.

Benthic foraminifera incorporate Cd and Ca into their shells in

proportion to the ratio of these elements in seawater, which

allows for the reconstruction of deepwater Cd (and thus mac-

ronutrient) distributions.

8.16.2.2.1 Cd in seawaterCd in seawater has a nutrient-like profile, with depleted values

in warm surface waters and the most enriched concentrations

in intermediate and deep waters (Boyle et al., 1976). Many

high-quality Cd measurements in seawater over the last

20 years have confirmed this nutrient-like behavior, with Cd

showing a tight correlation with phosphorus in seawater.

While it is still not clear whether most Cd is taken up as an

essential micronutrient or inadvertently, the cycling within the

ocean is dominated by incorporation into organic tissue in the

surface ocean and regeneration of elemental Cd as this material

is respired below the surface. The relationship with phospho-

rus is not linear, with a lower slope at lower phosphorus

concentrations (Figure 3). This curvature may be due to pref-

erential uptake of Cd over P by phytoplankton, though not at a

fixed (‘Redfield’) ratio (Elderfield and Rickaby, 2000). As dis-

solved phosphorus is drawn down in the sunlit surface ocean,

the Cd/P ratio of seawater decreases. When high-Cd/P organic

matter is remineralized below the surface, the Cd/P of seawater

climbs again, and the net result is a curve whose shape depends

on the effective fractionation factor between organic matter

Cd/P and seawater Cd/P. Since different deep waters have

undergone different extents of remineralization, and may be

mixtures of water masses with different preformed Cd/P ratios,

there are real differences in the Cd/P relationships in different

0.0

1.0

0.8

0.6

0.4

0.2

1.2

0.0 1.0

Estimated seawater Cd (nmol kg−1)

Cd

rec

onst

ruct

ed fr

om fo

ram

inife

ra (n

mol

kg−1

)

0.80.60.40.2 1.2

Figure 4 Seawater Cd reconstructed from foraminifera using adepth-dependent empirical distribution coefficient versus estimatedbottom water Cd. Data are from Boyle (1988, 1992).

1.2

1

0.8

0.6

0.4

Cd

SW

(nm

olkg

−1)

0.2

00.0 2.41.60.8

PSW (μmol kg−1)3.2

1.4

Figure 3 Seawater Cd measurements from the modern ocean plottedversus PO4. The curvature in this relationship is modeled by Elderfieldand Rickaby (2000) (solid line) as resulting from the preferential uptakeof Cd into marine organic matter. Reproduced from Elderfield H andRickaby REM (2000) Oceanic Cd/P ratio and nutrient utilization in theglacial Southern Ocean. Nature 405: 305–310.


ocean regions (Elderfield and Rickaby, 2000), which may of

course change through time. Because Cd is also removed from

seawater via CdS precipitation in suboxic sediments, the global

seawater Cd inventory, and thus the global Cd/P relationship,

can change as the extent of suboxic sediments changes

(Rosenthal et al., 1995a,b; van Geen et al., 1995). However,

the residence time for Cd in seawater is sufficiently long, and

the expansion of deep-sea suboxia sufficiently small, that the

Cd concentration during the last glacial period is unlikely to

have been lower than today by more than 5% (Boyle and

Rosenthal, 1996). This allows us to make meaningful infer-

ences about the cycling of major nutrients via reconstructions

of the past Cd concentration in seawater.

8.16.2.2.2 Cd/Ca in foraminiferaBoyle (1981) showed that with careful attention to cleaning,

benthic foraminiferal shells reflect the concentration of some

trace metals (including Cd) in seawater, with a preferential

incorporation of Cd over Ca into foraminifera. The empirical

distribution coefficient (D, ratio of Cd/Ca in foraminifera to

Cd/Ca in seawater) was shown to be 2.9, based primarily on

deep-sea core tops (Boyle, 1988). More core-topmeasurements

at all water depths have subsequently shown that this distribu-

tion coefficient increases with depth, starting at �1.3 above

1 km and increasing to a maximum of �2.9 below 3 km

(Boyle, 1992). Recent measurements suggest that the shallow

water value may be as high as 1.7 and is not discernibly

temperature dependent (Bryan and Marchitto, 2010; March-

itto, 2004). Once the depth-dependent distribution coefficient

is applied, the reconstructed Cd from foraminiferal tests shows

a linear relationship to Cd estimated for the overlying water

(Figure 4). Boyle et al. (1995) showed that the aragonitic

foraminifer Hoeglundina elegans has a D of 1 that shows

no depth dependence. McCorkle et al. (1995) observed an

unexpected decrease in the apparent distribution coefficient

of Cd and other trace metals with depth between 2.5 and

4 km in the deep Pacific, which suggested dissolution-driven

loss of Cd from the foraminiferal calcite after burial. Elderfield

et al. (1996) favored the idea that the lower D in deep Pacific

waters is a calcification response to the low carbonate satura-

tion state, rather than representing postdepositional dis-

solution, a possibility acknowledged by McCorkle et al.

(1995) and supported by laboratory dissolution experiments

(Marchitto et al., 2000). There has been no systematic attempt

to correct paleoceanographic Cd reconstructions for this

apparent low carbonate saturation state effect, though Zn/Ca

(Marchitto et al., 2000) and B/Ca (Yu and Elderfield, 2007)

hold some promise in this regard.

To the extent that the d13C of infaunal benthic foraminifera

reflects the low d13C of pore waters, it would be expected that

infaunal Cd/Ca also reflects the interstitial rather than the

bottom water concentrations. However, while interstitial d13Cvalues decrease with increasing depth in the sediment due to

the regeneration of isotopically light organic matter, the distri-

bution of Cd in pore water is more complex. In shallow pore

waters, Cd increases due to regeneration of organic material

but in suboxic sediments Cd decreases due to precipitation of

CdS (McCorkle and Klinkhammer, 1991; Rosenthal et al.,

1995a,b). As a consequence, we would not necessarily expect

a systematic relationship between the Cd/Ca of epifaunal

versus infaunal foraminifera as there is for d13C. Tachikawaand Elderfield (2002) showed that for three sites in the North

Atlantic, pore waters were 2–4 times enriched in Cd relative to

bottom waters. They found that P. wuellerstorfi, whose d13Cvalues suggested that they calcified in bottom waters at the

sediment surface, had a D of 3–4, consistent with culture

studies (Havach et al., 2001). The infaunal species showed

Figure 4

Figure 3


a lower D of about 1 with respect to the pore water Cd con-

centrations. The net effect is that the Cd/Ca in P. wuellerstorfi

and the infaunal species is quite similar, consistent with previ-

ous observations across a range of conditions (e.g., Boyle,

1992). Further, Tachikawa and Elderfield argue that the

depth-dependent D for the infaunal Uvigerina may actually

reflect a trend in the enrichment of pore water Cd relative to

bottom water Cd, though this cannot explain the depth depen-

dence of D in the epifaunal P. wuellerstorfi (Boyle, 1992).

Regardless of the details surrounding infaunal D values, it is

expected that Cd/Ca in P. wuellerstorfi or other surface dwelling

foraminifera should give a more direct measure of bottom

water Cd. As with d13C, careful interpretation aided by consid-

eration of supplemental information about changes in bottom

water characteristics has yielded regionally coherent and sensi-

ble patterns of deepwater Cd (e.g., Boyle, 1992; Marchitto and

Broecker, 2006).

8.16.2.3 Ba/Ca

Like Cd, dissolved Ba in a broad sense cycles like a nutrient,

depleted in surface waters and higher in deep waters, but the

surface depletion is moderate, and the regeneration of Ba

occurs deeper than for labile nutrients like phosphorus. This

pattern is due to the removal of Ba by barite precipitation in

decaying organic matter and subsequent barite dissolution in

the deep sea. This process coincidently results in a close corre-

lation between Ba and alkalinity in today’s ocean (Lea, 1993).

Lea and Boyle (1989) showed that the barium content in

benthic foraminiferal shells is controlled by the barium con-

centration in bottom waters, with a D of �0.4. Ba/Ca in

foraminifera has been used, like Cd/Ca, as a paleotracer of

water masses (e.g., Lea and Boyle, 1990) and appears to suffer

from the same carbonate saturation state-linked effect as

Cd/Ca (McCorkle et al., 1995).

8.16.2.4 Zn/Ca

Zn is a necessary micronutrient for phytoplankton and espe-

cially diatom growth. This requirement results in a very strong

correlation between oceanic dissolved Zn and Si concentrations

(Bruland et al., 1978). Benthic foraminiferal Zn/Ca is controlled

by bottom water dissolved Zn concentration, with a maximum

D in P. wuellerstorfi and Uvigerina of �9 (Marchitto et al., 2000)

and a value closer to 22 in C. pachyderma (Bryan and Marchitto,

2010). Because of the close relationship between Si and Zn,

Zn/Ca in foraminifera is particularly useful for tracing the

Si-rich bottom water that originates in the Southern Ocean

(Marchitto et al., 2002). Like Cd/Ca and Ba/Ca, the apparent

distribution coefficient for Zn decreases in waters which are

more undersaturated with respect to calcium carbonate. Because

they respond to saturation state differently, Cd/Ca and Zn/Ca

can theoretically be used together to assess both the nutrient

content and carbonate ion concentrations in deep water masses

(Marchitto et al., 2000, 2002).

8.16.2.5 Silicon Isotopes

A promising new proxy for the reconstruction of deepwater

silicic acid concentrations is the isotopic composition of Si in

sponge spicules. Hendry et al. (2010) have shown that the

fractionation of Si isotopes increases with ambient Si(OH)4concentration. This effect is thought to be related to the growth

rate of the sponge spicules, which should respond to the

Si(OH)4 concentration.

8.16.3 Conservative Water Mass Tracers

In general, nutrient-type tracers have been used to infer the

sources and mixing of deep water masses in the past, despite

the complication that the nutrient content of deep water de-

pends both on the nutrient content at its source and the

addition of nutrients via regeneration of organic matter as the

water mass ages. The utility of this method is primarily due to

the large and readily measured contrasts in the nutrient-related

tracers. Headway is now being made in the use of conservative

tracers that reflect the mixing of source waters alone (such as

temperature and salinity in the modern ocean).

8.16.3.1 Mg/Ca in Benthic Foraminifera

Mg/Ca ratios in benthic foraminifera reflect the calcification

temperature (e.g., Martin et al., 2002; Rathburn and

deDeckker, 1997; Rosenthal et al., 1997b). In principle, tem-

perature can be used as a conservative tracer for subsurface

water masses, but at present, the uncertainties in converting

Mg/Ca to temperature overwhelm the deep water temperature

gradients that distinguish water masses of different origins. In

particular, the saturation state of bottom water appears to exert

a strong secondary influence on Mg/Ca, as it does for the other

trace metal ratios (Elderfield et al., 2006). Emerging research is

targeting additional ratios like Mg/Li (Bryan and Marchitto,

2008; Lear et al., 2010) and infaunal taxa like Uvigerina

(Elderfield et al., 2010) in the hope of isolating or minimizing

the carbonate ion effect so that Mg/Ca could be used to assess

the contributions to intermediate water masses, where the

temperature signals are more easily distinguished.

8.16.3.2 Pore Water Chemistry

The d18O of seawater is set in the short term by processes which

occur at the ocean surface (evaporation, precipitation, sea ice

formation and melt; Craig and Gordon, 1965), so it can be

used as a conservative tracer in the subsurface ocean. The d18Oof glacial-age water has been reconstructed by measuring the

d18O of pore water in deep-sea sediment (e.g., Schrag and

dePaolo, 1993; Schrag et al., 2002). The maximum contribu-

tion of high d18O from the LGM appears at a depth of 20–60 m

within the sediments. The actual d18O of the glacial-age deep

water must be reconstructed by modeling the diffusion of

water within the sediments. Estimates for the bottom water

temperature (also a conservative tracer) at the LGM can then be

made by using the benthic foraminiferal d18O along with

information from pore waters about the d18O of the water in

which it calcified. Similarly, salinity can be reconstructed using

pore water [Cl] (Adkins and Schrag, 2001). While the number

of locations with such data is limited, Adkins et al. (2002)

show the promise of using pore water d18O and [Cl] to


reconstruct the properties of deep water masses. As more data

are collected, these tracers will also be useful for reconstructing

the sources and mixing of deep water masses.

8.16.3.3 Oxygen Isotopes in Benthic Foraminifera

The d18O of foraminiferal calcite reflects the d18O of the water

in which it calcified and the temperature of calcification. Since

both temperature and d18O of seawater are imprinted on a

water mass at the surface, the d18O of foraminifera is a conser-

vative tracer that can be used to determine the sources and

mixing of water masses in the deep ocean. The use of d18O as a

conservative tracer does not require the separation of the tem-

perature and d18O/salinity components of the foraminiferal

d18O signal. Despite the fact that oxygen isotope measure-

ments are routinely gathered simultaneously with carbon iso-

topes, the use of d18O as a tracer for ocean circulation has

lagged behind the use of d13C. Zahn and Mix (1991) showed

that gradients in deep water foraminiferal d18O are small due

to the fact that NADW is warmer than Antarctic Bottom Water

(which would lead to lower foraminiferal d18O), while d18O of

NADW is higher (counteracting the temperature effect). They

also demonstrated that there are significant issues with inter-

laboratory and interspecies calibration that interfere with the

effective use of this tracer. However, recent studies have shown

that Cibicidoides and Planulina do a good job of tracking calci-

fication temperature (Figure 5; Duplessy et al., 2002; Lynch-

Stieglitz et al., 1999a). While early studies showed that Cibici-

doides d18O did not reflect equilibrium calcification (O’Neil

et al., 1969), later determinations of equilibrium calcification

are consistent with the observed values for core-top Cibicidoides

(Kim and O’Neil, 1997; Lynch-Stieglitz et al., 1999a). Some of

the interlaboratory calibration problems arise because most

mass spectrometry labs routinely calibrate using a carbonate

standard (NBS-19) that can be up to 6% lower in d18O than

benthic foraminifera from the deep ocean during glacial

3

2

1

0

0Temp

δ18O

calc

ite (P

DB

) − δ

18O

wat

er (‘

PD

B’)

105

−1

−2

−3

4

Figure 5 The isotopic fractionation, expressed as the difference between thSMOW to a PDB-like (Peedee Belemnite) scale by subtracting 0.27, in accordcalcification. Cibicidoides and Planulina from Lynch-Stieglitz et al. (1999a) (sDuplessy et al. (2002) (circles, solid line is a linear regression through theseexperiments of Kim and O’Neil (1997) (triangles).

periods. Care must be taken to calibrate regularly with stan-

dards spanning a large range of isotopic ratios and appropriate

corrections must be made for many mass spectrometers (e.g.,

Ostermann and Curry, 2000). While the range of foraminiferal

d18O in the deep ocean is small, the foraminifera seem to be able

to record these differences well. Mackensen et al. (2001) and

Matsumoto et al. (2001) also showed regionally coherent pat-

terns of benthic foraminiferal d18O in the Southern Ocean.

Lund et al. (2011) used vertical profiles of benthic foraminiferal

d18O to constrain the relative strength of transport and mixing

in the deep glacial Atlantic. With careful attention to laboratory

calibration, deep and intermediate water foraminiferal d18O re-

constructions show great promise as a conservative tracer for

ocean circulation.

8.16.3.4 Clumped Isotopes

There has been much recent progress in developing clumped-

isotope thermometry in carbonates. The enrichment of 13C–18O

bonds in carbonates relative to that expected from a random

distribution of these isotopes of carbon and oxygen (D47) ap-

pears to be strongly temperature dependent (Ghosh et al.,

2006). This relationship appears to hold for both the calcite in

foraminifera (Tripati et al., 2010) and the aragonite in deep-sea

corals (Thiagarajan et al., 2011). A thermodynamically based

paleothermometer from oceanic carbonates which would allow

for both temperature reconstructions and, when combined with

d18O of the calcite, water column d18O reconstruction consti-

tutes a major development. However, the precision of this

method is insufficient at present to distinguish the temperatures

of deep water masses, and more work is needed.

8.16.3.5 Carbon Isotope Air–Sea Exchange Signature

Because the d13C in seawater is influenced by both biological

cycling and air–sea exchange, one can use data on co-occurring

252015erature (°C)

30

e d18O of foraminiferal calcite and the d18O of seawater (converted fromance with paleoceanographic convention) versus temperature ofmall pluses, dashed line is a linear regression through these data) anddata) are shown along with results from the inorganic precipitation

Figure 5


nutrients (which are influenced only by biological cycling) to

isolate d13C differences that arise from air–sea exchange

(Lynch-Stieglitz and Fairbanks, 1994; Oppo and Horowitz,

2000; Rickaby and Elderfield, 2005; Zahn and Stuber, 2002).

Because the air–sea exchange signature is reset only at the

surface, it is a conservative tracer for ocean circulation. Today,

newly formed NADW and Weddell Sea Bottom Water have

d13C air–sea exchange signatures that differ by 0.8% (Lynch-

Stieglitz et al., 1995). However, the patterns in the air–sea

exchange signature are only as good as the contributing d13Cand Cd/Ca data, and areas where either dissolution or produc-

tivity overprints may be a problem must be avoided, limiting

the use of this tracer.

8.16.4 Neodymium Isotope Ratios

Deep water masses of the ocean have distinctive neodymium

isotope ratios that are recorded in manganese nodules

(Albarede and Goldstein, 1992; Albarede et al., 1997), dis-

persed manganese–iron oxides in deep-sea sediment (e.g.,

Gutjahr et al., 2007; Rutberg et al., 2000), and iron oxides on

fish teeth (Martin and Haley, 2000) and in foraminifera (e.g.,

Klevenz et al., 2008; Vance et al., 2004). Seawater is thought to

acquire its Nd isotope signal at the ocean boundaries, via

surface inputs (dust and riverine) and exchange with sediment

(Frank, 2002; Goldstein and Hemming, 2003; Lacan and

Jeandel, 2005). Away from the boundaries, especially in the

deep Atlantic where the residence time of water is short

compared to the residence time of Nd, the isotopes of this

element can serve as a quasiconservative tracer (Piepgras and

Wasserburg, 1980). Nd isotope ratios in NADW are low,

reflecting the addition of Nd from continentally derived detri-

tal material, which is predominantly old, whereas Nd isotope

ratios in the North Pacific are high, reflecting input from recent

volcanic activity. In regions where there is active mixing be-

tween these water masses (the Atlantic Ocean and the Southern

Ocean), both seawater and the Fe–Mn oxyhydroxides that

precipitate from it have intermediate values, as is the case for

the nutrient and conservative water mass tracers.

8.16.5 Circulation Rate Tracers

The water mass tracers discussed above can be used to assess the

sources of subsurface water masses and the mixture of waters

from different sources. However, this information can be used

only to assess the relative rates of renewal of these deep water

masses from the various source regions (e.g., Legrand and

Wunsch, 1995). There are several geochemical techniques that

can be used to assess the rates of ocean circulation.

8.16.5.1 Radiocarbon

Carbon-14 (radiocarbon) is produced in the atmosphere by

cosmic rays and decays with a half-life of 5730 years. Natural

radiocarbon levels in the ocean are highest at the surface where

they partially equilibrate with atmosphere. Deepwater concen-

trations are highest in newly formed NADW and decrease to

the lowest values in the deep Pacific where the deep waters

have been out of contact with the atmosphere for longest

(Stuiver et al., 1983). Benthic foraminifera record the radiocar-

bon concentrations of the deep water in which they grew, but

the 14C in the foraminifera decays over time, with >99% gone

after 40000 years. In order to reconstruct the seawater radio-

carbon content at the time the foraminifera were living, radio-

carbon measurements can be made on benthic and planktonic

foraminifera from the same interval in the core (Broecker et al.,

1988, 1990; Shackleton et al., 1988) in order to determine the

age of the benthic foraminifera. However, the radiocarbon

content of the planktonic foraminifera will reflect not only

their age but also the initial radiocarbon content of the near-

surface water in which they grew. The initial radiocarbon in

surface and near-surface waters can be significantly less than

expected for equilibrium with the atmosphere due to the rela-

tively long time required for carbon isotopic equilibrium be-

tween the surface ocean and atmosphere. One must also

account for the fact that the radiocarbon content in the atmo-

sphere changes with time, a change which impacts the initial

radiocarbon content of surface waters as well. This means that

deep waters at a given locationmay have initially formed under

a different atmospheric radiocarbon burden than that recorded

by codeposited planktonic foraminifera (Adkins and Boyle,

1997).

Also complicating the use of concurrent benthic and

planktonic radiocarbon concentrations to assess water mass

radiocarbon is the necessary assumption that benthic and

planktonic foraminifera from the same depth interval are

actually the same age. If the peak abundances of planktonic

and benthic foraminifera occur at different depths in a core,

bioturbation can introduce a spread in the planktonic–benthic

ages in cores with low accumulation rates, which is unrelated

to changes in bottom water age (Broecker et al., 1999; Peng

et al., 1977). While these problems can be overcome by using

high-sedimentation-rate cores (e.g., Keigwin and Schlegel,

2002), the widespread use of this technique has been limited

primarily by the identification of suitable cores with high

enough foraminiferal abundances for accurate radiocarbon

determinations. Sediments with high accumulation rates tend

either to have low foraminiferal abundances (due to dilution

by terrigenous material) or to be in high-productivity regions

of the ocean where surface waters are expected to be far from

equilibrium with the atmosphere and variable through time.

Another approach to constraining the age required to ex-

tract seawater 14C content from benthic foraminifera is to use

stratigraphic correlation. Numerous cores with high accum-

ulation rates record seawater characteristics such as surface

temperature or oxygen content that vary in concert with either

Greenland or Antarctic air temperatures, as recorded by ice

cores. Assuming the oceanic changes are nearly synchronous

with the polar records, the ice core chronologies (which are

independent of 14C) can be applied to the sediments and

therefore to the benthic foraminifera (Skinner and Shackleton,

2004). This approach is mainly limited by the quality of the

correlation and the assumption of synchroneity.

A third way to measure the deepwater radiocarbon ventila-

tion age in the past is by using deep-dwelling benthic corals

(Adkins et al., 1998; Goldstein et al., 2001; Mangini et al.,

1998). Here, the age of the coral can be independently deter-

mined using uranium series dating, allowing the radiocarbon


content of the coral to be used to determine the radiocarbon

content of deep water at the time the coral grew. This method is

unaffected by bioturbation but is currently limited by the

availability of deep-sea benthic corals.

Changes in the ventilation of the ocean can also be assessed

by looking at how the 14C content of the atmosphere has

changed through time (e.g., Hughen et al., 1998). The atmo-

spheric 14C content is controlled both by the rate of 14C pro-

duction in the atmosphere from cosmic ray bombardment and

the rate at which 14C is transferred into the ocean. While a

decrease in the ventilation of the ocean causes a buildup of

radiocarbon in the atmosphere, one must also account for

changes in production (believed to be dominated by changes

in geomagnetic field strength) before interpreting the atmo-

spheric radiocarbon changes in terms of ocean circulation

(Hughen et al., 1998; Lal and Charles, 2007).

8.16.5.2 231Pa/230Th Ratio

Both protactinium and thorium are produced in the water

column and scavenged by particles onto the seafloor. Because

protactinium is scavenged less efficiently by marine particles

(Anderson et al., 1983), much of the protactinium produced

in the Atlantic Ocean is transported to the Southern Ocean by

deep waters before it is scavenged onto the seafloor. This

produces a low (below production) 231Pa/230Th ratio in

Atlantic Ocean sediments and a high 231Pa/230Th ratio in

Southern Ocean sediments (Yu et al., 1996). Reconstruction

of the 231Pa/230Th ratio suggests that this pattern of Pa export

persisted in glacial times, and a low deepwater residence time

for the LGM Atlantic has been inferred (Yu et al., 1996). Time

series of Pa/Th from the North Atlantic have been used to infer

changes in circulation over the deglaciation (Gherardi et al.,

2009; McManus et al., 2004).

However, the 231Pa/230Th ratio in sediments depends on

the complex interplay between deep water movement and the

differential scavenging of the two elements throughout the

ocean basin (Marchal et al., 2001; Siddall et al., 2007). The

scavenging of Pa and Th depends both on the composition of

the particles and the particle flux. Biogenic opal fractionates Th

from Pa much less effectively than other particle types; as a

consequence, ratios in Southern Ocean sediments will be very

sensitive to particle composition (Luo and Ku, 1999; Yu et al.,

2001). While the Atlantic 231Pa/230Th fluxes are low and today

predominantly reflect the short residence time of deep waters

in this basin, in the Pacific ocean, 231Pa/230Th ratios are most

sensitive to particle flux, with higher ratios where the scaveng-

ing is most intense at the boundaries (Yu et al., 2001). Robust

paleocirculation reconstructions from 231Pa/230Th ratios will,

in general, require detailed knowledge of particle fluxes and

composition (Gil et al., 2009; Lippold et al., 2009) as well as

good spatial coverage from a number of sediment cores (Burke

et al., 2011; Siddall et al., 2007).

8.16.5.3 Geostrophic Shear Estimates from d18O in BenthicForaminifera

Physical oceanographers routinely use the distribution of

density in the ocean to help quantify ocean circulation. A

balance between the pressure gradient force and Coriolis

term (geostrophic balance) is assumed. Velocity can then be

calculated from density information using either a measured or

assumed velocity at a reference level or by combining the

geostrophic balance with other constraints, such as conserva-

tion of tracer flow across a section. This approach has also been

applied to paleoceanography by using vertical density profiles

reconstructed from the oxygen isotopic composition of fora-

miniferal calcite (Lynch-Stieglitz et al., 1999a).

The d18O from calcite tests of benthic foraminifera preserved

in ocean sediments can be used to estimate upper ocean density

because both the d18O of calcite (d18Ocalcite) and seawater den-

sity increase with increasing salinity or decreasing temperature

(Lynch-Stieglitz et al., 1999a). For times in the geologic past, the

ability to reconstruct density from d18Ocalcite is most limited by

knowledge of the relationships between d18Owater and salinity as

well as the relationships between temperature and salinity. How-

ever, it is the gradients in density that are used for geostrophic

calculations, and the reconstructed density gradients are less

affected by these limitations than the density itself.

This technique has been used to reconstruct the flow of the

Gulf Stream in the Florida Straits by reconstructing vertical

density profiles on either side of this current (Lund et al.,

2006; Lynch-Stieglitz et al., 1999b, 2009, 2011). Perhaps more

useful for reconstruction of the large-scale ocean circulation

is the fact that the density difference between the eastern and

western margins of the ocean basins reflects the strength of

the geostrophic overturning circulation (Lynch-Stieglitz, 2001;

Lynch-Stieglitz et al., 2006).

At present, the vertical density structure for the past ocean

can be reconstructed using benthic foraminifera only where the

seafloor intersects the upper water column, at ocean margins,

islands, and shallow seamounts. Although planktonic forami-

nifera calcify at various depths within the upper water column,

we have no way to reconstruct quantitatively the depth at which

they calcified. However, the oxygen isotopic composition of

deep-dwelling planktonic foraminifera can be used to recon-

struct the spatial pattern and direction of upper ocean flows

(Bostock et al., 2006; LeGrande and Lynch-Stieglitz, 2007;

Matsumoto and Lynch-Stieglitz, 2003; Ortiz et al., 1997).

8.16.6 Nongeochemical Tracers of Past OceanCirculation

While this is a geochemistry volume, any review of tracers of past

ocean circulationwould be incomplete withoutmention of some

of the sedimentological andpaleontological approaches to recon-

structing past ocean flow. Sediment grain size has been used

to reconstruct current intensity since the 1970s (Ledbetter and

Johnson, 1976). More recent grain size studies emphasize the

mean size of the ‘sortable silt’ fraction (>10 mm), which is least

likely to display cohesive behavior (McCave et al., 1995). Stron-

ger currents show larger sortable silt size due to changes in

both deposition and winnowing with current speed. The mag-

netic properties of sediments (concentration of magnetic

particles) have been used to infer flow paths and strength of

the overflows in the North Atlantic (Ballini et al., 2006; Kissel

et al., 1999).

Benthic foraminiferal assemblages have been related to dis-

tinct deep water masses and have been used for tracking these


water masses during the past (Schnitker, 1974; Streeter, 1973).

In more recent years, it has been determined that the deep-sea

benthic assemblage responds strongly to the flux of organic

material from overlying surface productivity (e.g., Loubere,

1991; Miller and Lohmann, 1982), limiting its use as a water

mass tracer to low-productivity regions (Schnitker, 1994).

0

(a)

(b)

Dep

th (m

)

(c)

1000

2000

3000

4000

5000

6000

0

0

1000

2000

3000

4000

5000

1000

2000

3000

4000

5000

40°S EQ20°S

40°S EQ20°S

40°S EQ20°S

Figure 6 (a) Modern distribution of the major nutrient PO4 in the western AtlanMackensen, 2004; Curry and Oppo, 2005), and (c) Cd (Marchitto and Broecker, 2Adkins JF, Curry WB, et al. (2007) Atlantic meridional overturning circulation duriAmerican Association for the Advancement of Science.

8.16.7 Ocean Circulation during the LGM

Some of these methods have been applied to examine ocean

circulation throughout the Cenozoic era, and others to exam-

ine century-scale changes during the very recent past. However,

all of the methods have been used to look at ocean circulation

1

1.5

2

2

1.5

1

0.5

0

−0.5

−1

−1.5

0.2

0.4

0.6

0.8

1

2.5

3

3.540°N20°N 60°N

40°N20°N 60°N

40°N20°N 60°N

tic (Conkright et al., 2002) and LGM distributions of (b) d13C (Bickert and006) inferred from benthic foraminifera. Reproduced from Lynch-Stieglitz J,ng the Last Glacial Maximum. Science 316: 66–69, with permission from the

Figure 6


during the LGM (about 19000–23000 years ago) and thus

have their maximum geographic coverage during this time.

Here, the picture of ocean circulation that these techniques

have produced for the LGM is briefly examined. The various

paleoceanographic methods outlined above each have their

advantages and pitfalls, times and places where they should

work well, and environments where variables other than ocean

circulation complicate the interpretations. There is no single

ocean circulation scenario for the LGM that is consistent with

all of the available data from all of the tracers. In the following

description, regionally coherent signals are weighted more

heavily than data from isolated sediment cores. Data in regions

where the complications may overwhelm the signal (e.g., high-

productivity regions for d13C data, the low-carbonate very deep

ocean for Cd data) are weighted less heavily. A quantitative

description of past ocean circulation will require many more

data and better understanding of the tracers, combined with

better techniques for dealing with the kinds of uncertainties

and multiple processes that affect the paleoceanographic

proxies.

Past circulation in the Atlantic Ocean basin has been

reviewed by Lynch-Stieglitz et al. (2007). The most prominent

feature of the glacial Atlantic is a water mass divide at about

2-km water depth. Both Cd/Ca and d13C distributions suggest

a strong stratification in the North Atlantic Ocean, with a low

nutrient, high-d13C water mass (often called Glacial North

Atlantic Intermediate Water, GNAIW) occupying depths

down to about 2000 m and a high nutrient, low-d13C water

mass of southern origin underneath (Figure 6; e.g., Boyle and

Keigwin, 1987; Curry and Oppo, 2005; Duplessy et al., 1988;

Marchitto and Broecker, 2006; Oppo and Lehman, 1993). An

expansion of southern-source waters into the deep Atlantic is

supported by higher Nd isotope ratios (Roberts et al., 2010)

40°W2000

1500

1000

500

0−2 0−1

26

27

25.5

26.5

27.5

sθ (kg m3)d18O

0°20°W

Holocene

20°E

(a) (b)

Dep

th (m

)

Figure 7 (a) Density of seawater across the South Atlantic at 30�S (Conkrigbasin, reflecting the presence of the meridional overturning circulation. (b) Oxrecent sediment on the eastern (red symbols) and western (blue symbols) sibasin today, with higher d18O for the higher density waters on the eastern siddata. (c) Oxygen isotopic composition of the shells of benthic foraminifera foin the upper 2 km was absent or reversed. Reproduced from Lynch-Stieglitz Jduring the Last Glacial Maximum. Science 316: 66–69, with permission from

and lower radiocarbon concentrations (Keigwin, 2004;

Robinson et al., 2005). The existence of these two distinct

water masses despite the mixing processes that would tend to

blend them suggests that one or both were renewed relatively

rapidly (Curry and Oppo, 2005), although reduced mixing

might have played a role as well (Lund et al., 2011).

It also seems clear at this point that the LGM 231Pa/230Th

ratio in Atlantic sediments was comparable to those we find

today (Gherardi et al., 2009; McManus et al., 2004; Yu et al.,

1996). This too would argue against a much slower circulation

of deep waters in the Atlantic Ocean during glacial times. The

cross basin contrast in d18O of benthic foraminifera, however,

was much smaller (Figure 7; Lynch-Stieglitz et al., 2006). This

would suggest that the glacial circulation was not simply a

shallower version of today’s meridional overturning circula-

tion, whereby surface and intermediate waters travel into the

North Atlantic to compensate the export of NADW.

Less is known about the state of the LGM Pacific Ocean,

largely because poor carbonate preservation in the Pacific

Ocean limits the availability of cores with benthic foraminif-

era. Undersaturation with respect to calcium carbonate in

the Pacific also complicates the interpretation of the trace

metal-based methods that suffer from carbonate ion effects.

However, as in the Atlantic, the carbon and oxygen isotope

data suggest a deep hydrographic boundary at about 2000-m

water depth (Herguera et al., 1992, 2010; Keigwin, 1998;

Matsumoto et al., 2002; Mix et al., 1991), with a low-d13C,high-d18O water mass below this divide and a high- d13C, low-d18O water mass above this divide. Like today, the lowest d13Cwaters in the Pacific are in the far north but at deeper depths

(>2000-m water depth) than today (Figure 8; Herguera et al.,

2010; Matsumoto et al., 2002). There is now considerable

evidence that intermediate waters in the North Pacific were

43210321 5

2000

1500

1000

500

0

(‰ PDB) d18O (‰PDB)

Glacial(c)

Dep

th (m

)

ht et al., 2002). The waters are denser on the eastern side of theygen isotopic composition of the shells of benthic foraminifera (d18O) inde of the basin. These profiles reflect the density contrast across thee. The solid lines are the predicted d18O based on modern hydrographicr the LGM. These profiles suggest that the modern density contrast, Adkins JF, Curry WB (2007) Atlantic meridional overturning circulationthe American Association for the Advancement of Science.

Figure 7

1

2

3

4

5

0.50.5 0.5

0.3 0.30.3

0

0.4−0.2

0.3 ± 0.1(n=5)

0.3

xxx xx

xxxxxx xxxxxx

xxxx

x

xx

xxxx

xxx

x?

xxxxx

xx xx x

xx

x

x

xxx

x

xxxxxx

x

xxxxx

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xxxxx

0.3

0.40.40.4

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−0.1 −0.1−0.1

−0.1

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?

0.3±0.2

0.4

(n=6)

0.3 ± 0.1(n=3)

0.1 ± 0.0(n=3)

−0.4±0.1(n=4)

0.2 ± 0.0(n=3)

0.0 ± 0.1(n=5)

−0.1 ± 0.1(n=5)

−0.1 ± 0.0(n=3)

−0.3 ± 0.1(n=4)

−0.4 ± 0.1(n=6)

−0.3 ± 0.1(n=5)

0.2±0.1(n=8)

0.2 ± 0.1(n=8)

−0.4 ± 0.1(n=4)

0.1 ± 0.2

Pacific Holocene d13C

Pacific LGM d13C

Wat

er d

epth

(km

)

1

2

3

4

5

80

South Latitude North

40200204060 60

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er d

epth

(km

)−0.2

−0.2

−0.1

−0.1

0.1

0.1

0.10.3

NPIW

(n=3)

Figure 8 Modern and LGM distributions of benthic foraminiferal (Cibicidoides) d13C in the Pacific Ocean. Reproduced from Matsumoto K,Oba T, Lynch-Stieglitz J, and Yamamoto H (2002) Interior hydrography and circulation of the glacial Pacific Ocean. Quaternary Science Reviews21: 1693–1704. With permission from Elsevier.


Figure 8


more oxygen-rich during the LGM (van Geen et al., 2003),

consistent with the d13C-based scenario of waters more de-

pleted in nutrients at intermediate depths. The LGM Indian

Ocean is also characterized by a benthic divide at about

2000 m, with higher d13C and lower Cd reconstructed for

the upper water mass (Boyle et al., 1995; Kallel et al., 1988;

McCorkle et al., 1998). This deep hydrographic boundary is

also seen off southern Australia (Lynch-Stieglitz et al., 1996).

For both the upper and lower water masses, there is a

general progression from highest d13C in the North Atlantic

to intermediate values in the South Atlantic and South

Australia to progressively lower d13C values in the North Pacific

(excluding the deepest Southern Ocean). Goldstein et al.

(2001) also argue, based on limited glacial-age radiocarbon

data, that the modern progression of youngest deep waters in

the North Atlantic, intermediate values in the Southern Ocean,

and oldest values in the deep Pacific was preserved during

glacial times. The d18O values in the upper water mass are

similar from North Atlantic to North Pacific, within the con-

siderable scatter in the data. Whether or not the GNAIW

reached the Pacific or there was another source of nutrient-

poor upper deep water forming in the North Pacific remains a

source of debate (Keigwin, 1998; Lynch-Stieglitz et al., 1996;

Matsumoto et al., 2002; Oppo and Horowitz, 2000), although

box model results do argue for GNAIW export out of the

Atlantic (Sigman et al., 2003).

Finally, there appears to be an even denser water mass that

occupies the deepest portion of the Southern Ocean. This deep

water has high d18O values and very low d13C, with lowest d13Cvalues (about �1%) observed in the Atlantic sector of the

Southern Ocean (Figure 6). While there was some thought

that the low d13C values may reflect a productivity-related

overprint (Mackensen et al., 1993), a convincing case has

been made that the d13C values are regionally coherent and

34.5

3

2.5

2

1.5

1

0.5

0

−0.5

−1

−1.5

−2

3.5

35.535

Modern North Atlantic

Modern South Pacific

Thet

a (°

C)

Modern Southern Ocean

Site 9812184 m N. Atla

Site 10634584 m N. Atlantic

Figure 9 Estimates of LGM potential temperature (theta) and salinity fromsalinities (assuming 125 m of eustatic seawater change) are indicated with thfor the LGM are close to the freezing point of seawater (�1.98 �C). The data tReproduced from Adkins JF, McIntyre K, and Schrag DP (2002) The salinity, twith permission from the American Association for the Advancement of Scie

unrelated to other measures of overlying productivity

(Mackensen et al., 2001; Ninnemann and Charles, 2002).

Nd isotope values within this water mass have less of a

North Atlantic signature than today (Rutberg et al., 2000).

Unfortunately, there is not enough spatial coverage at present

to document convincingly the extent of this deep water mass in

the Indian and Pacific sectors of the Southern Ocean. The Cd/

Ca data are even more limited than the d13C data in the

Southern Ocean, due largely to low foraminiferal abundance

and potential overprints from saturation state; these data none-

theless suggest that the low d13C values are not accompanied

by a significant increase in nutrients (Rosenthal et al., 1997a).

This suggests that reduced air–sea exchange may have been

responsible, at least in part, for the lower d13C in the deep

Southern Ocean relative to the rest of the deep LGM ocean

(Marchitto and Broecker, 2006; Toggweiler, 1999).

Reconstructions of salinity from pore water [Cl] (Adkins

et al., 2002) suggest that this deep Southern Ocean water

mass was considerably saltier than deep waters in the Atlantic

and Pacific basins (Figure 9). Adkins et al. (2002) also argue

that pore water d18O suggests that this dense, salty water is

produced as a by-product of sea ice formation. Furthermore,

there is limited evidence that this water mass had very low

radiocarbon levels, indicating isolation from the atmosphere

and very sluggish ventilation rates (Skinner et al., 2010). It is

therefore thought that this deep southern water mass was able

to store significant amounts of carbon and thus contribute

to lower glacial atmospheric CO2 levels (e.g., Sigman and

Boyle, 2000; Toggweiler, 1999). While this capacity would

have been greatly enhanced if the water mass filled all ocean

basins up to the �2-km LGM nutricline, the most extreme char-

acteristics (lowest d13C, highest salinity) appear limited to the

Southern Ocean, suggesting other deep water sources in the

world ocean.

3736.536 37.5

Last Glacial Max hydrography

nticSite 11233290 m S. Pacific

Site 10933626 m Southern O.

pore water profiles (Adkins et al., 2002). The modern and LGM meane thin and bold arrows, respectively. All deepwater temperature estimateshus indicate a cold, high-salinity water mass in the deep Southern Ocean.emperature, and d18O of the glacial deep ocean. Science 298: 1769–1773,nce.

Figure 9


8.16.8 Conclusions

Considerable strides have been made in the last twenty-five

years in developing methods to reconstruct past ocean circula-

tion. There are now a number of different geochemical

methods that can yield information on paleo-ocean flow.

However, as the use of each method develops, it becomes

clear that there are multiple processes affecting both the tracers

themselves and their imprint into the sediments. We are still

just beginning to understand most of these methods and there

are limited amounts of data available. A full understanding of

these factors and how to separate out their effects, combined

with multiple measurements using methods affected by differ-

ent processes, can ultimately lead to well-constrained recon-

structions of past ocean circulation. However, even now, these

methods are capable of giving us a tantalizing glimpse of the

ocean during the LGM, when the water masses were arranged

in a distinctly different pattern than they are today.

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