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Water. Flower and Kennett (1995) suggested that with the cessa-

tion of Tethyan outflow around 14 million years ago, Southern

Component Water production increased, thereby reducing the

meridional heat transport and making an expansion of the East

Antarctic Ice Sheet possible. Hence, reconstructions of the pro-

duction rates of warm and cold deep-water masses suggest that

the middle Miocene cooling and growth of the East Antarctic

Ice Sheet were controlled by changes in ocean circulation.

The latest Miocene, during the Messinian stage (7.12–5.32

million years ago, Berggren et al., 1995) was a time of major

climatic and paleoceanographic change. Large anhydrite/gyp-

sum deposits from the Messinian are found around the Mediter-

ranean region, suggesting that the Mediterranean Sea was

desiccated during this time period. During the late Miocene,

the connections between the Atlantic Ocean and Tethyan

Ocean (a precursor to the Mediterranean) were through the

Rifian Corridor in Morocco and the Iberian Portal in Spain.

These passages allowed the inflow of Atlantic waters in to

the Paleo-Mediterranean to replenish evaporative loss in this

basin. With the progressive closure of the two straits that con-

nected the Paleo-Mediterranean with the Atlantic, the Paleo-

Mediterranean became isolated and desiccated (e.g., Benson

et al., 1991; Hodell et al., 1994). This led to what is called

the Messinian salinity crisis. The desiccation of the Paleo-

Mediterranean affected oceanic salinity by removing an esti-

mated 6% of dissolved salts from seawater. The cause of the

Mediterranean salinity crisis was probably the interplay

between tectonic changes, glacio-eustatic changes and climate

changes within the Paleo-Mediterranean.

The Pliocene

The early Pliocene (5–3 Ma) seems to be the most recent time

period of sustained global warmth. During the mid-Pliocene

(3 Ma), prior to the onset of Northern Hemisphere glaciation,

global mean temperatures may have been as much as 3.5



C war-

mer than at present (Raymo et al., 1996). Much evidence comes

from the mid-Pliocene, when high-latitude sea surface tempera-

ture increases suggest strong thermohaline circulation.

The reasons for mid-Pliocene warmth are not yet clear.

Crowley (1996) summarized several possible reasons: chan-

ging CO

levels, orographic effects related to late Cenozoic

uplift and/or changes in meridional heat transport due to, for

example, closure of the Central American Isthmus. Recent stu-

dies call on a combination of two main factors to explain Plio-

cene warmth: increased atmospheric CO

levels and enhanced

thermohaline circulation. Because CO

is a “greenhouse” gas,

changes in Pliocene atmospheric CO

levels could be the cause

of the mid-Pliocene warmth. Raymo et al. (1996), for example,

suggested that atmospheric CO

levels during the mid-Pliocene

were on average about 35% higher than the pre-industrial

value. CO

estimates for the mid-Pliocene come from different

sources. Among these are studies of the carbon isotopic

composition of marine organic matter (Raymo et al., 1996),

stomatol densities in leaves (Van der Burgh et al., 1993) and

the carbon isotope difference between surface and deep waters

(Shackleton et al., 1992). However, the results from these dif-

ferent techniques show little increase in pCO

for the Pliocene.

In addition to changes in atmospheric CO

levels, change in

ocean heat transport, i.e., thermohaline circulation is one of the

major mechanisms used to explain past time intervals of greater

warmth, like that of the mid-Pliocene. The thermohaline circu-

lation plays a significant role in regulating global climate by

redistributing heat and moisture and influencing the rate of

exchange between the ocean surface and deep-ocean. At

present, the thermohaline circulation is composed of two

major water masses: North Atlantic Deep Water (NADW) and

Antarctic Bottom Water (AABW). NADW is produced in the

northern North Atlantic and spreads towards the south, influen-

cing other ocean basins. AABW, which is mainly produced in

the Weddell Sea and Ross Sea, spreads towards the low lati-

tudes in all ocean basins. The overturning cells of NADW

and AABW are interrelated, which means that the relative

Figure N2 Trends in global climate change during the Neogene. The data shown are based on deep-sea stable oxygen isotope records (d

from more than 40 Deep Sea Drilling Project and Ocean Drilling Program sites (Zachos et al., 2001). The stable oxygen isotope composition of

bottom-dwelling foraminifers reflects changes in global ice-volume and /or changes in deep-water temperature. Solid black line corresponds to

heavily smoothed values of raw data (grey crosses). Increases in d

O values correspond to climatic deterioration, i.e., larger global ice-volume

and / or cooler deep-waters (after Zachos et al., 2001).

610 NEOGENE CLIMATES

abundance of these water masses in ocean basins can be mod-

ified by a change in production rate of only one of them. How-

ever, there are problems associated with holding changes in

ocean heat transport responsible for greater warmth (Crowley,

1996). One of them is the need to identify the factors responsi-

ble for changes in ocean circulation. Another is that increased

heat transport would decrease low-latitude sea-surface tempera-

tures (SST). Cooler tropical SSTs during past warm periods

are matters of controversy. Raymo et al. (1996) presented evidence

suggesting that the mid-Pliocene (3 Ma) warmth was related to

increased NADW production, in addition to slightly increased

levels. Although increased NADW production might be

responsible for increased warmth in the North Atlantic, it is

not clear how this would result in global warmth. Changes in ther-

mohaline circulation primarily redistribute heat over the planet.

The initiation of NADW production occurred during the

Neogene. Data from deep-sea cores suggests that NADW was

being formed during parts of the middle Miocene (17–11

million years ago) and that the flux of NADW increased in

the late Miocene (11–5 million years ago), approaching mod-

ern values in the early Pliocene (5–3 million years ago).

There are two major hypotheses for the cause of the initiation

of NADW. One is that the shoaling and subsequent closure of

the Central American Seaway allowed the North Atlantic to

become saline enough for the formation of NADW (Maier-

Reimer et al., 1990). The other is that changes in the flux

of NADW are related to changes in the sill depth of the

Greenland-Scotland Ridge (Wright and Miller, 1996). Based

on modeling studies, Mikolajewicz and Crowley (1997) sug-

gested that the initiation of NADW might have been related

to the closing of the Central American Seaway, but that the

depth of the Greenland-Scotland Ridge controlled its outflow.

Paleoceanographic studies in the western equatorial Atlantic,

the eastern equatorial Pacific and in the Caribbean indicate that

the closure of the seaway between North and South America

resulted in a distinct reorganization of ocean circulation starting

at about 4.6 million years ago (Haug and Tiedemann, 1998).

The emergence of the Panamanian Isthmus intensified the Gulf

Stream and introduced warm, saline water masses to the high

northern latitudes. Evaporative cooling of this warm, saline

water mass would result in the formation of NADW. However,

Kim and Crowley (2000) suggested, based on an ocean general

circulation model, that the increased NADW production could

be a consequence of Pliocene warming rather than a primary

cause. Their results indicate that a warming over the Southern

Ocean produces lower rates of sea ice formation and southern

ocean deep-water outflow. This in turn, leads to an increase in

NADW outflow in the Southern Ocean and an increase in pole-

ward heat transport in the North Atlantic. Kim and Crowley

further postulated that due to the greater sea ice area in the

Southern Ocean, this region would be more sensitive to the

inferred slightly higher CO

levels during the mid-Pliocene.

Onset of North ern Hemisphere Glaciation

The initiation of major NHG seems to have been the culmina-

tion of a longer-term high-latitude Cenozoic cooling that began

in the early Eocene and is marked by the first indications of ice

sheets in Antarctica about 36 million years ago. The long-term

cooling brought the climate system of Earth to a critical state

for ice-sheet buildup in the Northern Hemisphere. The glacia-

tion of Greenland and the Arctic is believed to have begun in

the late Miocene. Small amounts of ice-rafted debris (IRD)

are detected in Nordic Sea sediment cores back to about 11

million years ago (Jansen et al., 2000 and references therein).

The timing of the onset of NHG varies to some extent bet-

ween different regions. Although the initiation of major glacial

intervals probably occurred close to the middle/late Miocene

boundary, major glacial events are not apparent in the Nordic

seas until just before 3 million years ago. An intensification

of glacial conditions is recorded at about 3.3 million years

ago. The onset of large-scale NHG is dated at approximately

2.75 million years ago on the Vøring Plateau and 3.3 million

years ago on the Iceland Plateau (Fronval and Jansen, 1996;

Jansen et al., 2000). The different timing probably reflects dif-

ferent responses of the Greenland Ice Sheet versus ice-sheets in

northern Europe. IRD with an origin from the Laurentide Ice

Sheet seem to appear around 2.74 million years ago. In general,

enrichment of the stable oxygen isotope composition of benthic

foraminifers suggests that there was significant deep-water

cooling and an increase in global ice volume after 3.2 Ma

(Tiedeman et al., 1994; Shackleton et al., 1995)(Figure N2).

Several different causes for the onset of NHG have been

suggested. One theory attributes the onset of NHG to a decline

in the concentration of atmospheric CO

. This would reduce

the amount of heat trapped in the atmosphere and lead to cool-

ing. Changes in atmospheric CO

could arise from increased

productivity or from a longer-term decline in CO

due to the

burial of organic carbon and chemical weathering. It has also

been hypothesized that the final closing of the South American

Seaway was linked to the onset of NHG. The strengthened

northward flow of warm, salty water would suppress sea-ice

formation. With a reduced sea-ice cover, more moisture from

the ocean would be available to the surrounding landmasses

and trigger the growth of large ice sheets there. Haug and

Tiedemann (1998), for example, proposed that increased atmo-

spheric moisture supply was a necessary precondition for ice-

sheet growth, which was then triggered by changes in the

Earth’s orbital obliquity. A contrasting argument is that

increased deep-water formation could have worked against

the initiation of NHG as greater heat transport to high latitudes

would have tended to prevent ice-sheet formation.

The geological record shows strong evidence that the ther-

mohaline circulation in the early Pliocene weakened as a

response to the onset of Northern Hemisphere Glaciation

(e.g., Raymo et al., 1992, 1996). Based on the stable carbon

isotope composition in benthic foraminifera from the Atlantic

and Pacific oceans, Raymo et al. (1992) associated the gradual

cooling between 3 and 2 Ma with an increased suppression

of the formation of NADW. Data from the eastern Pacific

(Shackleton et al., 1995) shows an in-phase relationship

between colder periods and those having decreased NADW

formation during the Pliocene. It has been argued that tectonic

changes such as the closing of the Central American Seaway,

uplift of the Himalayan and Tibetan Plateaus, and deepening

of the Bering Strait are too gradual to account entirely for the

rapid onset of NHG. Although tectonic changes may have

brought the global climate to a critical threshold, the rapid var-

iations in Earth’s orbital parameters and thus insolation are

more likely to have triggered the intensification of NHG (e.g.,

see Lourens and Hilgen, 1997). Long-term periodic variations

in the Earth’s orbital parameters seem at least partially to

have controlled the onset of NHG (see Astronomical theory

of climate change).

Carin Andersson

NEOGENE CLIMATES 611

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Cross-references

Antarctic Bottom Water and Climate Change

Antarctic Glaciation History

Astronomical Theory of Climate Change

Carbon Isotopes, Stable

Cenozoic Climate Change

Ice-Rafted Debris (IRD)

Marine Biogenic Sediments

Messinian Salinity Crisis

Mid-Pliocene Warming

North Atlantic Deep Water and Climate Change

Ocean Paleocirculation

Oxygen Isotopes

Plate Tectonics and Climate Change

Pre-Quaternary Milankovitch Cycles and Climate Variability

Thermohaline Circulation

Weathering and Climate

NITROGEN ISOTOPES

Introduction

Nitrogen, element number five in the Periodic Table, is a limit-

ing nutrient to primary production (i.e., plant growth) on the

continents and in the ocean (Dugdale and Goering, 1967). Phy-

toplankton productivity is generally high where fixed (i.e.,

soluble) nitrogen concentrations are elevated and low where

they are diminished. Because primary productivity influences

climate, via the uptake of atmospheric carbon dioxide, and is

the precursor to petroleum deposits, it is important to under-

stand nitrogen cycling in the environment. The stable isotopes

of nitrogen,

N and

N, representing, respectively, 99.63%

and 0.37% of nitrogen atoms in the solar system, provide

a powerful tool for elucidating the processes that transfer nitro-

gen between reservoirs because many of them cause large

isotopic fractionations that can be measured with a mass

spectrometer.

612 NITROGEN ISOTOPES

The global nitrogen cycle

The largest reservoir of nitrogen on the planet is rocks, which con-

tain 1.9  10

Tg (10

g) of nitrogen (Wada and Hattori, 1990).

Nitrogen in rocks is not very reactive and therefore has a very long

residence time. The largest available pool of nitrogen is atmo-

spheric dinitrogen gas (N

), consisting of 3.9  10

Tg N (Wada

and Hattori, 1990). For comparison, the total amount of nitrogen

contained in all plants and animals on the continents is 1.5  10

Tg N and in the ocean is 4.7  10

Tg N, both very much smaller

than the reservoir of nitrogen in the atmosphere. While some

specialized microbes are able to convert N

gas into ammonia

(NH

) – the biologically useful form of nitrogen – by a process

called nitrogen fixation, most primary producers require fixed

nitrogen in the form of nitrate NO





or ammonium ion



. The pool of nitrate in the ocean is 5.7 10

Tg N and that

of ammonium ion is 7  10

Tg N (Wada and Hattori, 1990).

The marine nitrogen cycle is complex and not well under-

stood (Codispoti, 1995). Current estimates indicate that the

total of all oceanic sinks for nitrogen are larger than combined

sources by 75% (Codispoti and Christensen, 1985)butuncertain-

ties of a factor of two to four exist regarding the magnitude of

each. Sources of fixed nitrogen to the ocean are (a) nitrogen

fixation (30–130 Tg N yr

–1

), (b) atmospheric deposition (30–83

Tg N yr

–1

), and (c) river runoff (13–24 Tg N yr

–1

) (Wada and

Hattori, 1990). Sinks of nitrogen in the ocean are (a) microbial

denitrification, the conversion of nitrate to N

gas with its subse-

quent loss to the atmosphere (25–180 Tg N yr

–1

), (b) burial in

marine sediments (20 Tg N yr

–1

), and (c) exports of organic nitro-

gen to the continents via fish catches, guano deposition,

and atmospheric transport (10–20 Tg N yr

–1

) (Codispoti and

Christensen, 1985).

While the combined exports of nitrogen substantially exceed

combined imports in the above tally, recent studies indicate that

nitrogen fixation rates may have been drastically underesti-

mated by a factor of two or more (Gruber and Sarmiento,

1997; Karl et al., 1997; Zehr et al., 2001). So the marine nitro-

gen cycle may be approximately in balance. Nevertheless

future studies are needed to better constrain the sources and

sinks of nitrogen in the ocean.

Nitrogen isotope ratios

Nitrogen isotopic ratios are typically measured on gaseous nitrogen

or N

O) in a dual-inlet mass spectrometer. Very high precision,

on the order of 0.01% or better, is obtainable using this technique.

The isotope ratio of a sample is reported relative to that of a refer-

ence standardand reported in delta notation in units of per mil (%),

where one per mil is one-tenth of one percent, or one part per

thousand, according to Equation (1):

N ¼



sample





standard



 1  1; 000

ð1Þ

Atmospheric nitrogen (N

) is the reference standard for

nitrogen isotopic analyses and has a d

N value defined as 0%.

The biologically-mediated reduction reactions that con-

vert nitrogen from nitrate NO



; þ5 oxidation state



to nitrite



; þ3



to nitrous oxide (N

O, + 1), to nitrogen gas

,0),andtoammonia(NH

, –3 ) ar e faster for

Nthanfor

N as a result of higher vibrational frequency of bonding to

N than to

N(Owens,1987). This results in products that are

N-depleted (i.e., have a lower

Nratioandlowerd

N) rela-

tive to the substrate. If the substrate reservoir is either closed or has

inputs and outputs that are slow relative to one of the reduction

processes then the reservoir will become enriched in

N. An

example is the oxygen-depleted sub-surface water of the eastern

tropical Pacific Ocean where high rates of denitrification cause

large

N enrichments of the nitrate pool on the order of 1% (or

10%) (Cline and Kaplan, 1975). Denitrification in low-oxygen

regions of the ocean, such as the eastern tropical Pacific and the

Arabian Sea cause the d

N value of the ocean nitrate reservoir

to be high (5%) relative to the atmosphere and to terrestrial

nitrogen (Wada et al., 1975; Sigman et al., 1999).

During the reduction of N

to NH

by nitrogen fixing bac-

teria in the ocean, atmospheric nitrogen (d

N=0%), which

is already isotopically-depleted relative to ocean fixed nitrogen,

is further depleted, resulting in organic nitrogen with a d

value of –2.7% (c.f., Sachs and Repeta, 1999 and references

therein). In regions of the ocean where nitrogen fixation is

extensive, such as in the subtropical north Pacific (Liu et al.,

1996; Karl et al., 1997) and in the Mediterranean Sea (Sachs

and Repeta, 1999), d

N values are low in both phytoplankton

and fixed nitrogen relative to global ocean average values.

Oxidation reactions that convert ammonia to N

O; NO



and



during nitrification also result in isotopic fractionation

(Mariotti et al., 1981; Yoshida et al., 1989).

Diagenetic alteration of nitrogen isotope ratios

Early diagenetic (i.e., decomposition) reactions in marine sedi-

ments can severely alter the isotopic composition of sedimentary

organic nitrogen. This is critical to consider when attempting the

interpretation of nitrogen isotopic ratios in whole sediment

because the effect of diagenesis can be as large as the primary

signal that is sought. For example, the diagenetic overprint

of sedimentary nitrogen isotopic values typically ranges from

3–8% in sediments overlain by well-oxygenated water (Altabet

and Francois, 1994; Sachs and Repeta, 1999), while the total

range of isotopic variation is typically less than 5%. The pro-

blem of diagenetic alteration of sedimentary d

Nvaluesisless

of a concern in sediments overlain by oxygen-depleted water

(Altabet et al., 1999; Sachs and Repeta, 1999).

One means of circumventing the issue of diagenesis is to

measure nitrogen isotopic ratios on molecular fossils, or “bio-

markers.” These molecules have a known source and, if found

intact, are unlikely to have been altered isotopically. Chloro-

phyll and its decomposition products, chlorins, are ideal for this

purpose because they are produced by all primary producers

and persist in sediments for millions of years (Sachs and

Repeta, 1999). Algal culture studies indicate that chlorophyll

N values reflect those of the host alga with a constant isoto-

pic depletion of 5.1% (Sachs et al., 1999). Methods for purify-

ing chlorins for nitrogen (and carbon) isotopic analysis are

described in Sachs and Repeta ( 2000 ).

Conclusion

Nitrogen isotopic ratios provide a powerful tool for evaluating

processes within the nitrogen cycle and for reconstructing

changes in the cycling of nitrogen through time. Because multi-

ple nitrogen transformation reactions occur simultaneously in

any environment, with some leading to nitrogen isotopic

enrichment and some to isotopic depletion of the reservoir

NITROGEN ISOTOPES 613

of interest, it can be difficult to discern the most important

processes. Nitrogen isotopes are therefore most useful in envir-

onments where one or more transformation processes can be

ruled out.

Julian P. Sachs

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Organic Geochemical Proxies

Stable Isotope Analysis

NORTH ATLANTIC DEEP WATER AND CLIMATE

CHANGE

Modern North Atlantic deep water

Together, the atmosphere and oceans transport 5–6PW

(PW = 10

W) of heat away from the tropics toward each pole

in response to the latitudinal imbalance between incoming solar

and outgoing long-wave radiation. In the Atlantic Ocean, net

ocean heat transport is northward in both hemispheres, with a max-

imum of 1.3 PW across northern middle latitudes (Ganachaud

and W unsch, 2000). This transport is dominated by the meridional

overturning circulation, which transforms warm surface currents

(moving predominately northward) into cold deep waters (moving

predominately southward). The main product of this transforma-

tion is North Atlantic Deep Water (NADW), which forms pole-

ward of 50



N and fills much of the Atlantic between 1and

4 km depth belowsea surface. NADW formation is thus associated

with the release of a great deal of heat to the high northern latitude

atmosphere.

Modern NADWis ventilated in two main regions of wintertime

deep convection: the Greenland Sea (producing Greenland Sea

Deep Water) and the Labrador Sea (producing Labrador

Sea Water) (Figure N3). In general, three factors may contribute

to convective instability: cooling, increased salinity (lack of fresh-

water input, or active salinification due to sea ice formation), and

winds/storminess. Sea ice also leads to the formation of dense

Arctic shelf waters that entrain warm Atlantic waters. The result-

ing Arctic Intermediate/Deep Water mixes with Greenland Sea

Deep Water to form the dense overflows that pass southward on

each side of Iceland (Denmark Strait Overflow Water to the west,

and Iceland-Scotland Overflow Water to the east). Each overflow

Figure N3 Schematic diagram of modern NADW formation, after

McCartney and Talley (1984). Warm (>4



C) surface currents are shown

in black, and dense deep waters are dashed. The curled ends of the

warm currents indicate sinking. Also shown are cold surface currents

(open arrows). Labels denote the Labrador Sea (LS), Greenland Sea (GS),

and Norwegian Sea (NS).

614 NORTH ATLANTIC DEEP WATER AND CLIMATE CHANGE

amounts to roughly 3 Sverdrups (Sv = 10

1

), and several

additional Sv are entrained south of the sills as geostrophy concen-

trates the net flow into a western boundary current (Hansen and

Østerhus, 2000). Labrador Sea Water adds perhaps another 5 Sv,

yielding a total of 15  2 Sv of NADWexiting the North Atlantic

(Ganachaud and Wunsch, 2000). The overflows mainly contribute

to NADW properties below 2 km depth while the less-dense

Labrador Sea Water mainly affects the 1–2kmrange.

Hydrographic time series indicate that fluxes of the various

NADW components have varied slightly over recent decades

(Dickson et al., 1996; Hansen et al., 2001). Changing patterns

of surface salinity and winds may suppress deep convection

in one region while enhancing it in another. Analogous but lar-

ger scale changes are believed to have occurred in the distant

past, with dramatic impacts on climate.

Last Glacial Maximum

The importance of the North Atlantic’s wind-driven circulation in

explaining the climate of the last ice age was hypothesized by

James Croll well over a century ago, but the abyssal circulation

did not receive attention until much later. Weyl (1968)proposed

that Pleistocene glaciations were initiated by reduced NADW for-

mation and expanded sea ice coverage, with both being caused by

lower sea surface salinities in the North Atlantic. Early reconstruc-

tions of Last Glacial Maximum (LGM, 21 ka) North Atlantic

surface temperatures based on assemblages of planktonic organ-

isms suggested that polar waters and sea ice expanded southward,

as expected if the warm, upper branch of the meridional overturn-

ing circulation was restricted. Reduced NADW has since come to

be regarded as a response to (and amplifier of) glaciation, rather

than its ultimate cause (Rahmstorf, 2002).

Most of what is known about the past distribution of

NADW is based on paleonutrient proxies. Modern NADW is

low in dissolved nutrients like phosphate and nitrate because

it is formed from surface waters that have been stripped

of these constituents by photosynthetic activity. In contrast,

Atlantic deep waters of southern origin (Antarctic Bottom

Water (AABW) and Circumpolar Deep Water (CDW)) are

formed from waters that contain abundant nutrients. Deep

Atlantic nutrient distributions therefore reflect the spatial extent

of these water masses and give some indication of “aging” (via

the collection of nutrients remineralized from decomposing

organic matter that rains from above) and mixing processes.

Since there is no way to directly measure past phosphate

or nitrate distributions, nutrient proxies are relied upon for

paleoceanographic reconstructions. One such proxy is d

C, a

measure of the ratio between

C and

C in the dissolved inor-

ganic carbon (DIC) pool. The preferential uptake of

C during

marine photosynthesis leaves low-nutrient surface waters, such

as those that form NADW, enriched in d

C (Kroopnick, 1985)

(Figure N4). The d

C of DIC is recorded by the d

Cof

benthic foraminiferal calcite (Duplessy et al., 1984), offering

a way to reconstruct the distribution of NADW through time.

Another important nutrient proxy is dissolved Cd, which is

highly correlated with phosphate in the modern ocean. Some

phytoplankton use Cd as a micronutrient and others may incor-

porate it incidentally during the uptake of other nutrients. The

Cd concentration of seawater is reflected in the Cd/Ca ratio

of benthic foraminiferal calcite (Boyle, 1992), again offering

a method for tracing paleo-NADW. Discrepancies sometimes

exist between d

C and Cd/Ca, however, pointing to compli-

cating factors such as air-sea carbon exchange, benthic forami-

niferal microhabitat effects, and calcite undersaturation or

dissolution. Care must therefore be exercised when interpreting

changes recorded by just one proxy.

Benthic foraminiferal d

C and Cd/Ca provide a relatively

detailed picture of deep Atlantic water mass geometries during

the LGM. Below 2 km, both proxies indicate higher nutrient

concentrations during glacial times, consistent with reduced

NADW production (Boyle, 1992; Oppo and Lehman, 1993;

Sarnthein et al., 1994; Curry and Oppo, 2005; Marchitto and

Broecker, 2006). High d

C and low Cd/Ca values at inter-

mediate depths (above 2 km) suggest that NADW was par-

tially replaced by a shallower water mass dubbed Glacial

North Atlantic Intermediate Water (GNAIW). This apparent

shoaling of northern source deep waters allowed high-nutrient

Figure N4 Section of modern seawater d

C in the western Atlantic Ocean, after Kroopnick (1985) (reprinted with permission from Elsevier). High

values trace the southward flow of NADW, which mixes with low-d

C AABW and CDW. Dots indicate locations of samples, and numbers along

the top axis indicate GEOSECS station numbers.

NORTH ATLANTIC DEEP WATER AND CLIMATE CHANGE 615

southern source deep waters (AABW and/or CDW) to pene-

trate farther into the North Atlantic (Figure N5). Two additional

paleonutrient tracers, Ba/Ca and Zn/Ca, support this view (Lea

and Boyle, 1990; Marchitto et al., 2002).

Ocean circulation models suggest that reduced sea surface

salinities and greater sea ice cover in the glacial North Atlantic

could cause not only a shallower mode of deep water production

(GNAIW), but also a shift in its production site to south of Iceland

(Ganopolski et al., 1998). This change in deep convection sites

would reduce ocean heat transport to higher latitudes even if the

rate of deep water production was unchanged. The question of

rate cannot be addressed with paleonutrient data alone, since their

distributions depend on additional factors such as initial sea sur-

face (“preformed”) values, biogeochemical cycling, and water

mass mixing. One method that can potentially be used to assess

actual flow rates is the

C age difference between coexisting

benthic and planktonic foraminifera, which is an estimate of the

time elapsed since the deep water mass was last in contact with

the atmosphere (“ventilation age”). Mixing with old (poorly venti-

lated) AABW/CDW must be accounted for, however, to isolate

any true aging of northern source waters. Limited LGM measure-

ments are consistent with older waters in the deep (3–4km)

Atlantic (Robinson et al., 2005), but the impact of mixing has

not been well constrained. An additional complication for

benthic-planktonic ages arises from temporal variations in the

C content of the atmosphere, a problem that can be avoided

by obtaining paired

C and U-series dates on deep sea corals

(Robinson et al., 2005).

231

Pa/

230

Th patterns in marine sediments

have been used to argue that LGM GNAIW must have been

exportedto the Southern Ocean at a rate similar to today, otherwise

more Pa (which is less particle reactive than Th) would have been

scavenged out of the water column within the Atlantic (Yu et al.,

1996). The differential behavior of Pa and Th is dependent on

particle flux and composition, however, adding uncertainty to

rate estimations.

A promising method for reconstructing the rate of the meri-

dional overturning circulation is paleo-geostrophy. The north-

ward flowing limb of the North Atlantic overturning is

geostrophic, i.e., the flow is balanced by a horizontal pressure

gradient and the Coriolis force. The strength of the flow can

therefore be calculated from the magnitude of the pressure gra-

dient. Above some “level of no motion” (where isobars are hor-

izontal) the pressure gradient must be accompanied by lateral

changes in density. Because benthic foraminiferal d

O is influ-

enced by both temperature and salinity, it is a good recorder

of density in waters warmer than 5



C, allowing paleo-

geostrophic calculations to be made. Lynch-Stieglitz et al.

(1999) showed that Gulf Stream transport through the Florida

Straits was only 15–18 Sv during the LGM. Today the flow

is 30–32 Sv, with the wind-driven subtropical gyre account-

ing for 17 Sv and the rest being due to meridional over-

turning. Since it is unlikely that the gyre circulation was

slower during the LGM (models do not support weaker winds),

this suggests that the overturning was either very weak, or

was accomplished in a way that did not draw surface waters

northward through the Florida Straits.

Although benthic foraminiferal d

O has also been used to

infer deep ocean paleotemperatures, the influence of seawater

O (which is related to both salinity and global ice volume)

has been difficult to quantify. Fortunately, independent meth-

ods for separating temperature and salinity have recently

become available. Initial indications from benthic foraminiferal

Mg/Ca (a temperature proxy) are that the deep equatorial

Atlantic was 3–4



C colder during the LGM than during the

Holocene, consistent with a greater influence of AABW

(Martin et al., 2002). Semi-direct estimates of salinity and

seawater d

O can also be made by analyzing glacial-age

waters that have been partially preserved as sediment pore

waters. When combined with benthic foraminiferal d

O, these

measurements suggest that the LGM deep Southern Ocean was

much saltier and denser, but not discernibly colder, than the

deep North Atlantic (Adkins et al., 2002). This is very different

from the modern relationship (AABW colder and fresher than

NADW), indicating a major change in the thermohaline proper-

ties of the deep circulation. All deep waters appear to have

been formed close to the freezing point, with salinity dominat-

ing the deep stratification.

Dansgaard-Oeschger cycles and Heinrich events

During the last ice age, temperatures over Greenland and in the

North Atlantic repeatedly and abruptly switched between cold sta-

dials and warm interstadials, with a periodicity of approximately

1,500 years (Dansgaard et al., 1993). Broecker et al. (1985) sug-

gested that these Dansgaard-Oeschger (D-O) cycles could have

been caused by switches between “weak NADW” and “strong

NADW” states, driven by oscillations in North Atlantic salinity.

The first convincing evidence for a millennial-scale change in

NADW flow was during the Younger Dryas (13–11.6 ka), a

brief return to glacial-like conditions in the middle of the last

deglaciation. Benthic foraminiferal Cd/Ca and d

C imply a

sharply reduced presence of NADW in the deep western North

Atlantic at this time (Boyle and Keigwin, 1987). NADW suppres-

sion during the Younger Dryas may have been caused by an

input of Laurentide Ice Sheet meltwater into the North Atlantic

via the St. Lawrence River. Paleonutrient records (Keigwin

Figure N5 Vertical profiles of North Atlantic (>40



N) benthic

foraminiferal d

C for the Holocene (open circles) and LGM (filled circles)

(Oppo and Lehman, 1993; Marchitto et al., 2002). LGM data have been

shifted upward by 120 m to account for the glacial lowering of sea

level. Also shown are nearby modern seawater measurements (gray

line) (Kroopnick, 1985). The LGM nutricline at 2,000 m separates

high-d

C GNAIW from low-d

C AABW /CDW.

616 NORTH ATLANTIC DEEP WATER AND CLIMATE CHANGE

and Boyle, 1999; Boyle, 2000) and benthic foraminiferal faunas

(Rasmussen et al., 1997) now suggest that NADW production

was also reduced during D-O stadials and enhanced during inter-

stadials (Figure N6). There is evidence for an opposite response

(reduced during interstadials) for GNAIW (Zahn and Stüber,

2002), suggesting that stadials resembled the LGM state while

interstadials approached the Holocene state. Some deep waters

may have continued to form in the Norwegian-Greenland Seas

during stadials because of brines resulting from sea ice formation

(Dokken and Jansen, 1999).

Various ocean circulation models support the hypothesis that

even small freshwater forcing can abruptly switch NADW

between different equilibrium states (Rahmstorf, 2002). These

models exhibit hysteresis behavior, whereby the system flips

rapidly from strong to weak/no overturning, or vice versa, after

crossing some surface salinity threshold. Furthermore, the over-

turning can be bistable, meaning that it can exist in either

state for a given range of freshwater forcing. Glacial boundary

conditions may narrow the hysteresis loop, such that a smaller

range of salinity forcing is required to cause abrupt switches.

Periodic or quasi-periodic salinity “triggers” may result from

internal ice sheet dynamics (iceberg calving) (Schmittner et al.,

2002) or possibly solar variations. When combined with white-

noise salinity variability, the required magnitude of salinity for-

cing is further reduced, a phenomenon known as stochastic

resonance (Alley et al., 2001).

Some models’ glacial responses may be grouped into three

states: “warm” or strong NADW (interstadials), “cold” or weak

GNAIW (stadials), and “off” (Heinrich events) (Rahmstorf,

2002). Heinrich events, which occurred during stadials every

6–10 Kyr, are marked by widespread ice-rafted debris layers

believed to have resulted from ice sheet surging. The modeled

Heinrich state thus results from a much larger freshwater (iceberg)

forcing than in the normal stadial state. There is good paleonutrient

evidence for reduced NADW during Heinrich events (Boyle,

2000;Elliotetal.,2002), and atleast some GNAIW sites show

strong nutrient enrichments (Zahn and Stüber, 2002).

231

Pa/

230

measurements from Bermuda Rise suggest that the meridional

overturning circulation was nearly or completely shut down dur-

ing Heinrich event 1 (McManus et al., 2004).

Northern Hemisph ere glaciations (past 3 Myr)

Composite records of deep ocean d

O indicate that significant

Northern Hemisphere glaciations commenced during the late

Pliocene, at 3 Ma. Deep ocean d

C gradients suggest that a

gradual weakening of NADW began around this time (Ravelo

and Andreasen, 2000), probably as a response to (and amplifier

of) surface cooling and freshening. The divergence of Nd isoto-

pic values between the North Atlantic and Southern Ocean

since 3–4 Ma is consistent with this weakening (Frank et al.,

2002). On the 41 Kyr scale of late Pliocene glacial-interglacial

cycles, NADW was generally weakest during glacials and pro-

gressively deteriorated toward the Pleistocene. Glacial suppres-

sion of NADW intensified after 1.5 Ma and was most severe

during the 100 Kyr cycles of the mid-late Pleistocene (since

0.9 Ma) (Raymo et al., 1997). GNAIW has appeared during

glaciations (intermittently) since at least the early Pleistocene

(McIntyre et al., 1999).

On a finer temporal scale, there appear to be lags between

Pleistocene glacial cycles and NADW production changes

(Wright and Flower, 2002). During glacial initiations, NADW

initially strengthened as ice sheets began to grow, likely in

response to increased salinities. The resulting continued sea sur-

face warmth may have supplied excess moisture for the expand-

ing ice. At least several millennia passed before reduced salinities

caused NADW suppression and North Atlantic cooling. A simi-

lar (but opposite) progression seems to have occurred during

deglaciations: GNAIW initially weakened due to meltwater addi-

tion, with NADW resumption and surface warming occurring

several millennia later. Hence, NADW production changes have

played an important role in modifying the character of glacial-

interglacial transitions.

With still higher temporal resolution, it is apparent that the

millennial scale variability characteristic of the last glaciation

was also present during prior glaciations. Ice rafting events have

been associated with NADW reorganizations since at least the

early Pleistocene (Raymo et al., 1998). During the late Pleisto-

cene (past 0.5 Myr), millennial-scale sea surface variability was

strongest when ice sheets exceeded a certain threshold size

(McManus et al., 1999). Benthic d

C, however, suggests that

NADW reductions occurred during intervals of cooling and

ice rafting regardless of their magnitude, implying that rapid

deep circulation changes may occur without large Northern

Hemisphere ice sheets.

Earlier Cenozoic history

On longer timescales, tectonics played an important role in

the broad evolution of NADW. Prior to the global cooling of

Figure N6 Last ice age climatic and deep water indicators from

Bermuda Rise (4,600 m) in the western North Atlantic (Keigwin and

Boyle, 1999). Planktonic foraminiferal d

O shows warmings coincident

with Dansgaard-Oeschger interstadials (numbered). Benthic

foraminiferal d

C increased during interstadials, suggesting stronger

production of NADW. Cd /Ca measurements across one

stadial-interstadial transition are consistent with this (note reversed

scale). Low bulk percent CaCO

during stadials may have been due, in

part, to dissolution by corrosive AABW.

NORTH ATLANTIC DEEP WATER AND CLIMATE CHANGE 617

the Plio-Pleistocene, a relatively cold NADW mass (loosely

defined) could only form at very high latitudes, such as in the

Norwegian-Greenland Seas. During the early Cenozoic, a shal-

low Greenland-Scotland Ridge prevented any such waters from

filling the open Atlantic. Attempts to date the onset of NADW

formation, presumed to coincide with an opening in this barrier,

have relied on two main methods: reconstructing deep ocean

C gradients and identifying sedimentation patterns created

by deep flow. It has been proposed that an early Oligocene initia-

tion coincided with the formation of the Southeast Faroes Drift

(Davies et al., 2001). Accumulation on Feni Drift also began

around this time, though its link to NADW production is not

straightforward. The opening of Drake Passage (probably during

the Oligocene) may have also established the ocean dynamics

necessary to sustain a meridional overturning circulation, thereby

amplifying the cooling and glaciation of Antarctica (Toggweiler

and Bjornsson, 2000). The transition of deep Southern Ocean

C from low Pacific-like values toward high North Atlantic

values during the early-middle Miocene (20–15 Ma) has

been interpreted as the first appearance of substantial NADW

(Wright and Miller, 1996). The formation of other major North

Atlantic drifts around this time may or may not be related to

increased NADW flow.

Another tectonic change that affected NADW formation was

the closing of the Isthmus of Panama. Although the closure

began 13 Ma, the gateway was not shallow enough to

strongly impact NADW until 4.6 Ma (Haug and Tiedemann,

1998). By this time, the flow of relatively fresh Pacific waters

into the Atlantic was minimized, allowing the North Atlantic

to become saltier. This excess salt intensified the production

of NADW, which may have contributed to early Pliocene

warmth. The stronger northward flow of warm waters may

have also preconditioned the Northern Hemisphere for glacia-

tion at 3 Ma by supplying the requisite moisture for ice

growth.

Thomas M. Marchitto, Jr.

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Cross-references

Antarctic Bottom Water and Climate Change

Carbon Isotopes, Stable

Cenozoic Climate Change

Dansgaard-Oeschger Cycles

Heinrich Events

Last Glacial Maximum

Ocean Paleocirculation

Ocean Paleoproductivity

Oxygen Isotopes

Quaternary Climate Transitions and Cycles

Thermohaline Circulation

NORTH ATLANTIC OSCILLATION (NAO) RECORDS

Introduction

Natural climate variability in the Atlantic sector results from

processes at work in the atmosphere and in the ocean, and

interactions between the two. In general, Atlantic climate varia-

bility has been described as a function of three related compo-

nents: (a) an atmospheric mode of variability known as the

North Atlantic Oscillation (NAO), (b) an oceanic mode of

variability known as Atlantic Meridional Overturning Circula-

tion, and possibly (c) an atmosphere-ocean mode of dynamics

in the tropical Atlantic (see review by Marshall et al., 2001).

The first component, the NAO, is the leading mode of atmo-

spheric variability in the Atlantic sector and has been dubbed

the regional manifestation of the Arctic Oscillation (AO) –

the leading mode of large-scale atmospheric variability in the

Northern Hemisphere (NH). The NAO/AO is second only to

the El Niño-Southern Oscillation (ENSO) in terms of the range

of its climate impacts.

Impacts

The NAO has a profound impact on both marine and terrestrial

ecosystems (Stenseth et al., 2002) in the Atlantic sector and

surrounding continents, stretching from the Caribbean to the

Arctic and from the eastern seaboard of the United States to

Siberia. In addition to temperature and precipitation, impacts

of the NAO are seen in agricultural production (Souriau,

2001), yields from fisheries (Otterson and Stenseth, 2001),

fertility and demographic trends of land animals (Post and

Stenseth, 1998), snow depth (Beniston, 1997), water resource

management (Cullen et al., 2002), hydropower generation

(Visbeck et al., 2002 ), hurricanes (Landsea et al., 1999), and

wave heights (Kushnir and Wallace, 1989; Kushnir et al.,

1997). The NAO is also the major factor controlling air-sea

interactions over the Atlantic Ocean; modulating the intensity

and location of upwelling/downwelling centers of the Atlantic’s

Meridional Overturning Circulation(MOC).Furthermore,there

exists the possibility that the NAO has been influenced by

anthropogenically forced climate change. This has led to the

suggestion that a significant NAO-related weakening of the

MOC in global warming scenarios is – while contentious – a

possibility (Marshall et al, 2001).

Definitions

The NAO refers to a meridional redistribution of atmospheric

mass between the center of subtropical high surface pressure

located near the Azores, known as the Azores High (AH),

and the center of subpolar low surface pressure near Iceland,

known as the Icelandic Low (IL). There is no single way to

define the NAO. Techniques include one-point correlation

maps that rely on identifying regions of maximum negative

correlation (Wallace and Gutzler, 1981), principle component

analysis using eigenvectors of the covariance matrix (Barnston

and Livezey, 1987), and clustering algorithms that identify cli-

mate regimes corresponding to the peak in the probability den-

sity function (Cassou and Terray, 2001).

In general, the NAO is described as being in either a posi-

tive or a negative phase. Synchronous strengthening (positive

NAO state) and weakening (negative NAO state) of the NAO

centers-of-action have been shown to result in distinct,

dipole-like climate change patterns between western Green-

land/the Mediterranean and northern Europe/northeast US/

Scandinavia (Walker, 1924; Walker and Bliss, 1932; van Loon

and Rogers, 1978; Rogers and van Loon, 1979). A lower-

than-normal (higher-than-normal) IL and higher-than-normal

(lower-than-normal) AH result in an enhanced (reduced) pres-

sure gradient and a positive (negative) NAO. During this + NAO

(–NAO) phase, surface winds and wintertime storms moving

from west to east across the North Atlantic are stronger

(weaker) than usual. As a result, wetter and warmer (drier and

cooler) winter conditions occur in northern Europe, Scandinavia

and the east coast of the US, while cooler and drier (warmer and

wetter) winters occur in eastern Canada and Greenland.

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