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From: Nitrogen in the marine environment
Edited by: Douglas G. Capone
ISBN: 978-0-12-372522-6
© Copyright 2008 Elsevier B.V
The Netherlands
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C H A P T E R
3 4
Nitrogen in Past Marine Environments
Eric D. Galbraith, Daniel M. Sigman, Rebecca S. Robinson,
and Thomas. F. Pedersen
Contents
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1. Introduction
2. Potential Changes in the Marine Nitrogen Cycle During the Past
2 Million Years
2.1. Inventory changes
2.2. Internal stabilizing feedbacks on the marine nitrogen inventory
2.3. Nitrate distribution and the nutrient-rich surface regions
3. Recorders of the Past Nitrogen Cycle
3.1. Nitrogen isotope systematics
3.2. Ocean d15Nnitrate through time
3.3. Sedimentary nitrogen isotope records
4. Observations and Interpretations
4.1. Glacial–interglacial inventory changes
4.2. Glacial–interglacial changes in nutrient-rich regions
4.3. Connections
5. Perspectives on Progress
5.1. Looking backward
5.2. Looking forward
References
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1. Introduction
Many reservoirs of biologically important elements turn over so rapidly at the
Earth surface that modest differences between input and output could either remove
them completely or double their concentrations in the geologically short span of
thousands to millions of years. Despite this precarious situation, evidence indicates
that the Earth has been remarkably stable in its habitability, allowing the chain of
multicellular life to persist for more than a half billion years. Such continuity requires
a conspiracy of feedbacks within the Earth system that stabilizes the availability of
life’s key ingredients. Yet we know little about these feedbacks, how they develop,
or the constraints that they impose on the environment and life. Beyond these
Nitrogen in the Marine Environment
DOI: 10.1016/B978-0-12-372522-6.00034-7
#
2008 Elsevier Inc.
All rights reserved.
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fundamental scientific questions, an understanding of the regulatory mechanisms is
the key to predicting how the system will respond to anthropogenic climate change.
The marine nitrogen budget provides an important test case in the broader effort
to understand the stabilizing environmental feedbacks on the Earth surface. Because
both the principal source (N2 fixation) and sink (denitrification) of fixed nitrogen in
the oceans are mediated entirely by marine organisms, the marine nitrogen (N) cycle
is particularly well-suited to the development of biological homeostasis, providing
an important contrast in this regard to nutrients such as phosphorus, silicon, and
iron. However, it is difficult to quantify, on the basis of temporally and spatially
limited modern observations, the sensitivities of the N budget to environmental
parameters. Fortunately, geological and glaciological archives have recorded past
events that approximate large-scale experiments in which the oceanic N budget
responds to naturally imposed forcings.
Although some studies have explored the nitrogen cycle of the distant past (e.g.,
Fennel et al., 2005; Jenkyns et al., 2001; Kuypers et al., 2004; Rau et al., 1987;
Saltzman, 2005), we focus on the cycles of ice ages and interglacial periods that have
dominated the last 2 million years of Earth history. Several aspects of these cycles
make them well suited to this application. First, they are relatively well recorded in
sediment and ice core records. Second, the glacial and interglacial periods are each at
least three times longer than the residence time of oceanic fixed N (3 kyr,
(Codispoti et al., 2001)). Thus, a given glacial/interglacial transition (such as the
last deglaciation) approximates a discrete natural experiment, with the ocean N
budget beginning near steady state, responding to the climate-driven forcing and
then potentially reaching a new steady state condition. It has been suggested that the
marine N budget during the last glacial maximum was markedly different from today
(Altabet et al., 1995; Falkowski, 1997; Ganeshram et al., 1995), which may have
impacted atmospheric concentrations of the greenhouse gases CO2 (Broecker, 1982;
McElroy, 1983) and N2O (Gruber, 2004; Suthhof et al., 2001). By comparing the
modern nitrogen budget to this alternate state from the recent geological past, we
gain insight into the potential for future change.
While the budget of fixed N is unique among the nutrients, the internal cycling
of nitrogen and phosphorus appear to be tightly coupled through the stoichiometry
of algal uptake and subsequent remineralization in the modern ocean (Redfield,
1934). Thus, if we can understand the controls on the oceanic distribution of nitrate,
we would gain similar insight into the phosphate distribution. It has been posited
that coupled changes in polar ocean overturning and algal nutrient consumption
drive the observed variation in atmospheric CO2 over glacial/interglacial cycles (see
Sigman and Boyle (2000) for a review). If the relationship of high latitude nitrogen
cycling with climate over past glacial/interglacial cycles can be determined, it would
also shed light on the oceanic control of pCO2.
In this chapter, we review the ongoing efforts to use sediment and ice core
records to understand the dynamics of oceanic fixed N, focusing on recent glacial–
interglacial cycles. Research has, up to this point, followed the reductionist approach
of trying to reconstruct either changes in the ocean N budget (largely through low
latitude records) or changes in the internal cycling and distribution of N (largely
through high latitude records). However, evidence is building for a coupling of high
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and low latitude changes, connecting the processes that control the distribution of N
in the ocean with those that control the size of the whole ocean N inventory. We
begin by considering how the nitrogen cycle of times past may have differed from
that of today, followed by an overview of how nitrogen cycle features are recorded
in marine sediments and, finally, summarize the observations and interpretations
made from available records.
2. Potential Changes in the Marine Nitrogen Cycle
During the Past 2 Million Years
We divide the potential changes frequently considered by paleoceanographers
into two broad categories: those that most directly involve changes in the ocean
nitrogen inventory, that is, the total mass of fixed (non-gaseous) nitrogen in the
oceans, and those that most directly involve changes in the distribution and cycling
of fixed nitrogen within the ocean.
2.1. Inventory changes
Fixed nitrogen is continuously added to the ocean surface and removed from its
interior by the processes depicted schematically in Fig. 34.1. As discussed elsewhere
in this volume, the emerging view of the Holocene marine nitrogen budget is
largely independent of terrestrial nitrogen inputs and is dominated by the single
N2 fixation
P availability (+)
Fe availability (+)
Terrestrial N:
via rivers, atmosphere
interglacial
glacial
Sedimentary denitrification
Sea level (shelf area) (+)
Organic matter rain (+)
Oxygen supply (-)
Water-column
denitrification
Oxygen supply (-)
Organic matter rain (+)
Sedimentation of organic N
Figure 34.1 Schematic view of the major sources (green arrows) and sinks (white arrows) of
fixed N in the marine environment. Important factors for controlling the variation of each
source or sink term over time are given in italics, with (þ) and () indicating direct and indirect
proportionality, respectively.
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input term of marine N2 fixation and by the output terms of sedimentary and water
column denitrification (where ‘‘denitrification’’ refers to the total conversion of
fixed N to gaseous products by any pathway, including anammox). Given their
dominance with respect to mass flux, the potential for inventory changes is determined by the temporal variability of these three terms. As described below, it has
been suggested that the N inventory was larger during the last glacial maximum
(LGM), which would have allowed for greater biological sequestration of CO2 in
the deep sea, helping to explain the glacial pCO2 decrease.
2.1.1. Marine nitrogen fixation
Fixed nitrogen is supplied to the global ocean by specialized microbes, termed
diazotrophs, near the ocean surface (see Chapter 4, Carpenter and Capone this
volume). It has been proposed that the low supply rate of iron to the present-day
ocean hampers the growth rates of diazotrophs, exacerbating nitrogen limitation
among marine ecosystems (Falkowski, 1997). If the degree of global Fe-limitation
varies over geological time, then the rate of N2 fixation might be expected to follow.
It is believed that greater dust fluxes during glacial times (Petit et al., 1999; Winckler
et al., 2008) probably would have led to higher trace metal concentrations in the
oceans, though it is not clear by what degree (Parekh et al., 2004). If true, this could
have allowed diazotrophs to fix N at greater rates during ice ages, and the resulting
increase in N2 fixation rates could have increased the whole ocean N inventory
(Falkowski, 1997). Comparison of ice core records of atmospheric pCO2 and dust
led Broecker and Henderson (1998) to promote this idea.
2.1.2. Sedimentary denitrification
Organic matter is rapidly delivered to continental shelves of the modern ocean,
producing sediments that are generally anoxic very close to the sediment-water
interface. The physical proximity of oxygen-bearing bottom waters to anoxic
sediments with high concentrations of labile organic matter allows diffusion and
bioturbation to effectively supply oxygen to nitrifying organisms and nitrate to
nitrate-reducing organisms in the sediment, leading to high rates of denitrification
in shelf sediments (see Chapter 6, Devol, this volume). During ice ages, sea level
dropped by up to 120 m, exposing 75% of global shelf area (Hay and Southam,
1977). As a result, organic matter from highly productive coastal ecosystems would
have been shunted to the deep sea. Because this organic matter must settle through a
greater height of water column, more organic matter oxidation would have been
accomplished there instead of in sediments. One would expect, therefore, that the
lowering of sea level would have significantly reduced the global rate of sedimentary
denitrification, as proposed by Christensen et al. (1987). To obtain a rough maximal
estimate of this effect, we can assume that the area-normalized rates of sedimentary
denitrification on slopes and in the deep sea remained constant during glacial
maxima, despite the fact that they would be the primary depositional sites for
coastally produced sinking organic matter during low sea level. In this extreme
case, a reduction in global shelf area to 25% of the modern area would have
produced a decrease in sedimentary denitrification of 75 Tg N/year (Middelburg
et al., 1996), representing a 30% decrease in the global rate.
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Hypsometry is not the only parameter affecting global sedimentary denitrification rates. A higher flux of organic matter to the sediment strongly increases
sedimentary denitrification (see Section 2.2.2). The oxygen content of bottom waters
has a secondary and less certain effect. In some cases, higher denitrification rates
would be expected under more oxygen-poor conditions (Middelburg et al., 1996),
though this may be reversed if nitrification becomes limited by oxygen (Devol and
Christensen, 1993).
2.1.3. Water column denitrification
The other large sink of marine nitrogen is by denitrification in suboxic regions of
the water column. This process only occurs in the most poorly ventilated regions
of the modern ocean, where dissolved oxygen has been depleted through aerobic
respiration, to below 2–5 mM O2 ((Codispoti et al., 2005) see Chapter 6, Devol, this
volume). In the modern ocean, this is limited to less than 0.2% of the ocean volume,
almost entirely within the upper 1000 m of the Eastern Tropical Pacific and the
Arabian Sea. These subsurface regions are driven to suboxia by the coincidence of
rapid organic matter delivery with a slow re-supply of dissolved oxygen (Olson et al.,
1993; Sarma, 2002; Wyrtki, 1962). In theory, the extent of water column denitrification could be significantly altered by changing the export rate of organic matter to
the subsurface ( Altabet et al.,1995; Ganeshram et al., 2000), and/or by changing the
oxygen supply through a physical mechanism (Broecker and Peng, 1982; Galbraith
et al., 2004; Meissner et al., 2005). The relative balance between surface ocean
fertility and the oxidizing capacity of the subsurface can be approximated by the
ratio of nutrient to oxygen concentrations of global thermocline waters; all else
being equal, when this is high, denitrification is abundant. Because the concentrations of nutrients and oxygen in the thermocline are both strongly influenced by
those at high latitude outcrop regions (Galbraith et al., 2004; Meissner et al., 2005;
Sarmiento et al., 2004), water column denitrification has an acute sensitivity to the
surface conditions at high latitudes (see Section 4.3).
Since nearly all subsurface respiration on the global scale is supported by oxygen,
there is no clear a priori reason that water column denitrification should have been
similarly important at all times in the past; increases in the rate of oxygen supply
relative to consumption could have nearly eliminated it. However, the strong
sensitivities of water column denitrification, while rendering this process susceptible
to climate-driven variability, probably also make it more responsive to stabilizing
negative feedbacks (see Section 2.2.3).
2.2. Internal stabilizing feedbacks on the marine nitrogen inventory
The conceivable potential for nitrogen source and sink fluxes to vary independently
of each other would appear to characterize the marine nitrogen cycle as highly
unstable. However, modern process studies have illuminated stabilizing (negative)
feedback mechanisms that would tend to limit excursions in the N cycle (see Fig.
34.2). The strengths of these feedbacks on centennial to millennial timescales should
determine the responsiveness of the marine N cycle to climate change, but they
remain largely unquantified.
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Denitrification
N2 fixation
Subsurface
respiration
N inventory
N:P of
nutrient
supply
Productivity
Figure 34.2 Hypothesized feedbacks within the marine N cycle on centennial-millennial timescales. A direct proportionality, i.e., producing a change of equal sign to that of the forcing, is
shown as a black arrow, while an indirect proportionality is shown as a grey arrow. Note that each
complete loop contains exactly one indirect proportionality, therefore each represents a negative
(stabilizing) feedback on the fixed N inventory. Modified after Deutsch et al. (2004).
2.2.1. N2 fixation self-limiting feedback
It is widely held that N2 fixers tend to be at a competitive disadvantage to nondiazotrophs under nitrate-replete conditions: only where nitrate is nearly exhausted
but phosphorus remains available are N2 fixers expected to provide substantial new
nitrogen. If so, N2 fixation would tend to be self-limiting: greater rates increase new
N input to an ecosystem, reducing the degree of N limitation of non-diazotrophic
competitors and, hence, the net availability of phosphate to diazotrophs (Broecker,
1982; Redfield et al., 1963; Tyrrell, 1999; Deutsch et al., 2007). This self-limiting
feedback, illustrated on the left side of Fig. 34.2, intimately ties N2 fixers to the global
aggregate rate of denitrification, which sets the global deficiency of N relative to
P and, consequently, the availability of nitrate-limited environments exploitable by
N2 fixers. The sensitivity of this feedback depends largely on the degree of disadvantage that diazotrophs suffer in the presence of nitrate, which is poorly quantified.
The residence time of marine P is 3–15 times longer than that of fixed N (Delaney,
1998; Ruttenberg, 2003) and it has been argued that its source and sink fluxes are less
sensitive to glacial-interglacial climate change than those of fixed N (Ganeshram et al.,
2002). Given this apparent stability of the P inventory, the relatively robust relationship
between N and P concentrations in bulk algal stoichiometry (Redfield et al., 1963) is
critical when considering potential changes in the global N2 fixation rate. This ratio
reflects the nutritional requirements of the surface ocean algal community, thereby
determining its ability to utilize a given phosphate supply with available nitrate, and
hence, the potential for excess phosphate to be available to N2 fixers. Although there is
some degree of localized discrepancy from the canonical ‘‘Redfield’’ N:P value of 16:1,
including among individual species and in local environments, the ratio holds remarkably well throughout the modern ocean (see Gruber (2004) for a review and Chapter 1,
Gruber, this volume). Broecker and Henderson (1998) (Section 2.1.1) invoked a
global biomass N:P of 24:1 during glacial periods to explain lower atmospheric
pCO2. This degree of flexibility is not immediately suggested by the low variability of
remineralized N:P in the modern ocean, particularly with near-Redfield N:P in the
modern Fe-replete subtropical North Atlantic (Hansell et al., 2004). Nonetheless,
nitrate to phosphate ratios in the deep Mediterranean Sea can be as high as 25:1
(Krom et al., 2005), indicating that substantial divergences above Redfield N:P are
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possible under some conditions. Without a better appreciation of the mechanisms
underlying this master variable, the possibility of significant changes in the N:P of past
oceans cannot be dismissed, and the strength of the N2 fixation self-limiting feedback
remains in question.
2.2.2. Nitrogen-limitation feedback on denitrification
Denitrification in the ocean interior only occurs under extremely low oxygen
concentrations. Oxygen is depleted by the respiration of reduced carbon compounds, most of which are derived from the ecosystems above. If the ability of
those ecosystems to export fixed carbon is impaired due to nitrogen limitation of
export production, the resulting alleviation of the subsurface oxygen deficiency will
cause a reduction in denitrification. This constitutes a general stabilizing feedback on
both types of denitrification, illustrated on the right side of Fig. 34.2.
The extent of water column suboxia is sensitive to organic matter flux from the
surface, and would quickly respond to changes in export production. This sensitivity
is evident in the delicate balance between oxygen supply and oxidant demand in the
global subsurface: the respiration history of suboxic waters (oxygen utilization plus
denitrification) is almost entirely supported by oxygen, typically having consumed
250 mM O2 versus 15 mM NO3 (Codispoti et al., 2001). Denitrification in
suboxic waters is, therefore, a relatively small residual term in the difference between
total respiration and oxygen supply. As a result, small changes in the total respiration,
driven by export production, could have large effects on denitrification. Because
waters from denitrification zones supply the local upwellings that are, in turn, largely
responsible for maintaining the suboxia below, this is a geographically tight coupling
and would be expected to provide a strong feedback. The sensitivity of this feedback
depends on the degree to which diazotrophs near the upwelling zone are capable of
eliminating any N-limitation caused by a nitrate deficit in upwelled waters, thereby
mitigating changes in export production. In this regard, a stronger expression of the
N2 fixation self-limiting feedback posed above would weaken the denitrificationdriven feedback described here, and vice versa.
Sedimentary denitrification would respond similarly to a productivity change
caused by nitrate limitation, as it is largely determined by the organic carbon flux to
the seafloor (Middelburg et al., 1996). Hence, if a decrease in denitrification rates
promotes higher export production via a relaxation of nitrate limitation, the
enhanced organic carbon flux to the seafloor would reinvigorate sedimentary
denitrification. Conversely, an increase in denitrification encourages nitrogen limitation and limits large increases in export production. This reflexive feedback would
tend to dampen any massive swings in the N inventory.
2.3. Nitrate distribution and the nutrient-rich surface regions
The macronutrient-rich surface waters of the equatorial upwellings and polar oceans
are a particularly dynamic component of the global carbon cycle. In the Southern
Ocean, for example, the nutrient-rich and CO2-charged waters of the deep sea are
exposed to the atmosphere and returned to the subsurface before the available
nutrients are fully utilized by phytoplankton for carbon fixation. This incomplete
utilization of upwelled nutrients allows for the leakage of biologically sequestered
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CO2 back into the atmosphere, raising atmospheric pCO2 (Knox and McElroy,
1984; Sarmiento and Toggweiler, 1984; Siegenthaler and Wenk, 1984). On this
basis, increased nutrient utilization in the Southern Ocean has been invoked as the
cause of the lower pCO2 concentrations of glacial times, through a reduced leakage
of carbon into the atmosphere (Francois et al., 1997; Martin and Fitzwater, 1988).
The potential changes in the polar ocean over glacial/interglacial cycles are too
diverse to be usefully covered in this review (Sigman and Boyle, 2000; Sigman and
Haug, 2003). Nevertheless, it is illustrative to consider how the nutrient status of the
polar ocean would respond to a given change in ocean circulation, depending on the
parameters that most strongly control algal growth in these regions (a topic of
ongoing research in the modern ocean).
We consider the hypothesis that the polar oceans were more vertically stable
during the last ice age than they are today (Francois et al., 1997; Jaccard et al., 2005;
Robinson et al., 2004; Sigman et al., 2004; Toggweiler, 1999), focusing on the
Antarctic region. According to this hypothesis, increased density stratification of the
upper Antarctic water column would have reduced the rate of exchange between
nutrient-rich deep water and the euphotic zone. If algae growing in the Antarctic are
most critically light limited, their growth rates would be expected to have remained
similar regardless of climate state, so that decreased access to the deep nutrient pool
would have led to greater relative nutrient utilization (i.e. a greater ratio of nutrient
assimilation to gross nutrient supply).
On the other hand, if suggestions that Fe availability provides the overriding
control on algal production in the Antarctic are correct (Martin et al., 1991), the
predominant route of Fe supply to the surface would become a critical parameter. If
the Fe supply has always been dominated by dissolved Fe upwelled from below, its
supply would have decreased in step with the glacial decrease in the gross supply of
the macronutrients nitrate and phosphate (assuming a constant Fe: macronutrient
ratio in Antarctic subsurface water). In this case, the degree of relative nitrate
utilization in the glacial surface would have been more similar to that of today. By
contrast, if direct atmospheric Fe supply is important (either during the last ice age or
both during the last ice age and today), the relative utilization of nitrate and phosphate
in the glacial Antarctic surface would have increased with increasing stratification, as
in the case of light-limitation. Sedimentary records of export production, which
should be proportional to the net flux of macronutrients to the euphotic zone at
steady state, can be used to help untangle competing scenarios such as these. Clearly,
reconstructing the nutrient status of the polar oceans in the past and developing an
understanding of the controls on algal growth in these regions go hand in hand.
3. Recorders of the Past Nitrogen Cycle
Unfortunately, there are no existing direct records of past nitrate concentration; one must rely instead on proxy evidence. We focus here on the most specific
and frequently used proxy for studying the history of the ocean N cycle, the nitrogen
isotopic ratio of organic matter preserved in marine sediments.
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3.1. Nitrogen isotope systematics
There are two stable isotopes of N, 14N (which constitutes 99.63% of the N in the
environment) and the rare isotope 15N (see Sigman and Casciotti (2001) for a review
and Chapter 29, Montoya, this volume). Measurements of the d15N of nitrate
(d15Nnitrate) in the deep sea (below 1000 m) have shown it to be fairly homogeneous,
at about 5% relative to atmospheric N2 (Sigman et al., 2000) (d15N ¼ (15N/14Nsam14
15
ple/ N/ Nair –1) * 1000%). In certain regions and depth intervals of the ocean, the
15
d Nnitrate diverges significantly from the mean deep d15Nnitrate because of regionally
specific processes, as described below. To understand the signals that might be
retrieved from sedimentary N isotopes, we first look individually at each of the
key fractionating processes (Fig. 34.3).
3.1.1. Inputs
Inputs of nitrogen to the marine environment appear to be isotopically similar to
atmospheric N2. N2 fixation occurs with little effective isotope discrimination, so that
marine N2 fixers produce reduced nitrogen with a d15N only slightly lower than that
20
δ15Nnitrate (‰ vs. air)
15
Water column
denitrification
ε ~ 20‰
10
Nitrate
uptake
ε ~ 5‰
5
Sedimentary
denitrification
ε ~ 0‰
N2 fixation
δ15N ~ −1‰
0
0
0.5
1.0
1.5
2.0
Nitrate concentration (relative to initial value)
Figure 34.3 Nitrogen isotopic fractionation processes. The vectors approximate the isotopic
changes associated with the indicated addition or removal of nitrogen from an initial nitrate pool
with a d15N of 5% (black filled circle). Note that the fractionating processes (water column denitrification and nitrate uptake) are actually nonlinear, due to the progressive enrichment of d15N
in the residual pool, a feature not evident in this discrete view. The return of organic N to the
nitrate pool by remineralization is not shown despite a large isotope effect associated with nitrification (Casciotti et al., 2003), since ammonium in typical open-ocean subsurface water is
completely oxidized to nitrate. As a result, the nitrate produced should be similar to the organic
matter being remineralized. However, remineralization can lead to significant 15N enrichment in
the small fraction of organic matter that survives extensive degradation, as in the case of deep sea
sediments (see Section 3.3.1). After Sigman and Casciotti (2001).
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of air (0 to 2%, e.g., (Carpenter et al., 1997)). In regions of the subtropical ocean,
thermocline d15Nnitrate can be as low as 2%, 3% lower than the mean deep ocean
d15N. This observation is most readily explained by N2 fixation in the overlying
surface waters and the subsequent remineralization of this low-d15N organic material
in the waters below (Karl et al., 2002; Knapp et al., 2005; Liu et al., 1996; Pantoja et al.,
2002) (Fig. 34.4, North Atlantic).
Fixed N is also transferred from terrestrial to marine ecosystems by riverine and
atmospheric vectors, in roughly equal parts (Galloway et al., 2004). This probably
contributed no more than one quarter, and perhaps less than one eighth, of the input
flux of fixed N to the pre-industrial ocean (Galloway et al., 2004). Although it has
been suggested that the d15N of terrestrial inputs sums to near 0% (Brandes and
Devol, 2002), their isotopic compositions are very poorly constrained and may in fact
be distinct from the N introduced by oceanic N2 fixation. There is currently a need
for a systematic study of the isotopic composition of N inputs to the ocean from rivers
and groundwater that takes into account the multiple chemical forms of N involved
as well as the reactions that occur as rivers deposit N into the coastal ocean. Terrestrial
Eastern Tropical
North Pacific
Antarctic Zone,
Southern Ocean
δ15NO3− (‰ v. air)
δ15NO3−
6
8
10
12
14
4
North Atlantic
δ15NO3−
(‰ v. air)
5
6
7
0
1
2
(‰ v. air)
3
4
5
6
0
1000
δ15N
δ15N
[NO3−]
[NO3−]
[NO3−]
δ15N
Depth (m)
2000
3000
4000
5000
0
10
20
30
[NO3−] (μM)
40
0
10
20
30
[NO3−] (μM)
40
0
10
20
30
[NO3−] (μM)
40
Figure 34.4 Water column profiles of nitrate concentration (open symbols) and d15Nnitrate (filled
symbols) in the Eastern Tropical North Pacific (ETNP, coastal Baja California), Southern Ocean
and North Atlantic (Sargasso Sea).The ETNP shows a large increase in d15Nnitrate in the thermocline owing to local water column denitrification. The Southern Ocean shows little deviation
from the global deep mean d15Nnitrate except at the surface, where partial NO3 assimilation leaves
residual nitrate enriched in 15N. The North Atlantic profile shows low d15Nnitrate in the thermocline owing to the nitrification of locally fixed N. Note that the ETNP profile also includes deep
measurements from near Hawaii (diamonds); the smooth transition between samples at distant
locations emphasizes the homogeneity of the deep Pacific.
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inputs of fixed N are also a concern in paleoceanographic work given their potential
to be delivered directly to ocean margin sediments, thereby contaminating the
paleoceanographic signal in sediment records (see Section 3.3.2).
3.1.2. Outputs
Unlike N2 fixation, nitrate reduction has a strong isotopic bias, favoring the light N
isotope and leaving behind a residual nitrate pool that is progressively enriched in
15N with increasing degree of nitrate consumption. Culture studies collectively yield
a relatively broad range in the isotope effect, E (i.e. (14k/15k - 1)*1000%, where 14k
and 15k are the rate coefficients of the reactions for the 14N- and 15N-bearing forms
of NO3, respectively) for nitrate consumption by denitrifying bacteria, from
roughly 15% to 30% (Barford et al., 1999; Mariotti et al., 1981; Granger et al., in
press). By analogy with work on nitrate assimilation (Granger et al., 2004; Needoba
et al., 2004), it would appear that isotopic discrimination primarily takes place upon
reduction of nitrate to nitrite by the intracellular enzyme nitrate reductase which, in
its assimilatory form, has an intrinsic isotope effect of 15–30% (Ledgard et al., 1985;
Schmidt and Medina, 1991). The expression of this fractionation is registered by the
d15N of unconsumed nitrate that diffuses from the cells back into the ambient waters
(Granger et al., in press).
Water column studies, in which denitrification rates are calculated from nitrate
deficits relative to phosphate, have led to estimates of 22–30% for the isotope effect
of denitrification (Altabet et al., 1999b; Brandes et al., 1998; Cline and Kaplan,
1975; Liu and Kaplan, 1989; Sigman et al., 2003; Voss et al., 2001). Regardless of the
exact value of the expressed isotope effect, water column denitrification clearly
results in very high d15Nnitrate in the thermocline within, and adjacent to, actively
denitrifying regions of the water column (see Eastern Tropical North Pacific
(ETNP) profile, Fig. 34.4). We note that although there are no existing measurements to explicitly resolve the isotopic impact of anammox, it is implicitly included
in all existing water-column (and sedimentary) denitrification studies as part of the
overall fixed N loss.
Sedimentary denitrification, in contrast, shows very little expression of the enzymatic isotope effect. Benthic flux experiments have shown that, even with high rates
of sedimentary denitrification, minimal increase of d15Nnitrate is observed (Brandes
and Devol, 1997, 2002; Lehmann et al., 2004). The common explanation is that
nitrate is nearly completely consumed at the locus of sedimentary denitrification.
Because no nitrate escapes the denitrification zones, there is no expression of the
isotope effect. This view is supported by the observation of extreme 15N enrichment
of nitrate in sediment porewaters (Sigman et al., 2001; Lehmann et al., 2007), which
indicates that the lack of isotopic expression of denitrification in sediment cannot be
explained by a lack of isotopic fractionation at the scale of the organism. In any case,
in the context of budgetary calculations, sedimentary denitrification is typically
assumed to have a negligible isotopic effect (Brandes and Devol, 2002; Deutsch
et al., 2004; Sigman et al., 2003).
These two extremes of expression of the organism-level isotope effect of
denitrification—complete expression in the water column and no expression in
sediments—are unlikely to apply perfectly in the ocean. Instead, we might expect
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some degree of under-expression of the isotope effect in the water column under
some conditions (Deutsch et al., 2004; Sigman et al., 2003), particularly if alternate
dissimilatory pathways, such as anammox, fractionate differently and are of variable
importance depending on environmental conditions. An extreme case of this is
provided by the Cariaco Basin, where water column denitrification appears to
consume nitrate completely very near the oxic/anoxic interface, leading to minimal
expression of the isotope effect of denitrification (Thunell et al., 2004). In addition,
measurements of N2/Ar in the Arabian Sea (Devol et al., 2006) imply that the loss of
nitrogen through denitrification there is greater than previously thought, which
would suggest that the isotope effect of water column denitrification has been
significantly overestimated. Reciprocally, isotopic discrimination by denitrification
is partially expressed in some sediments, albeit to a small degree relative to water
column denitrification (Brandes and Devol, 2002; Lehmann et al., 2004, 2007).
Nevertheless, the end-member view presented above provides a useful conceptual
starting point.
3.1.3. Internal cycling
Fixed nitrogen occurs in many forms in the marine environment, and most transformations among nitrogen pools involve some degree of isotopic fractionation. The
most important fractionation for understanding changes in nitrate distribution (as
opposed to changes in the input/output budget) occurs during nitrate assimilation,
when discrimination by the enzyme nitrate reductase causes the d15N of the first
organic matter produced from a given pool of supplied nitrate to be lower than
d15Nnitrate. The d15N of both the residual nitrate and the subsequently produced
organic N increases as nitrate consumption progresses. Measurements of d15Nnitrate
in the upper ocean suggest an isotope effect of 4–10% for nitrate assimilation, with
most estimates closer to 5–8% (Altabet, 1988; Altabet and Francois, 2001; Altabet
et al., 1999b; DiFiore et al., 2006; Sigman et al., 1999b; Wu et al., 1999). Fractionation in cultures appears to be more variable, with isotope effects as high as 20%,
although many cultures yield effects of 6% (Needoba et al., 2003).
Developing a predictive understanding of the isotope effect of nitrate assimilation
has been a goal of marine N isotope studies. The isotope effect varies among species
grown under similar growth conditions (Needoba et al., 2003; Wada and Hattori,
1978; Waser et al., 1998) and, for at least some species, as a function of growth
conditions (Granger et al., 2004; Needoba and Harrison, 2004). For example, the
cultured diatom Thalassiosira weissflogii displays a greater organism-level isotope effect
under light-limitation than under iron-limitation or maximal growth rate (Needoba
and Harrison, 2004; Needoba et al., 2004). The potential variability of the isotope
effect represents a critical uncertainty in paleoceanographic application of nitrogen
isotopes to changes in nutrient utilization in the nitrate-replete regions of the surface
ocean.
When nitrate is completely utilized, the accumulated organic matter has the same
d15N as the original nitrate source. Hence, in regions where the consumption of
nitrate is complete over the course of the year, isotope fractionation during nitrate
assimilation has only a transitory (i.e. seasonal) effect that disappears when integrated
over the seasons of growth. From the perspective of paleoceanography, the d15N of
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organic N exported from such an environment must be identical to that of the N
supply; if subsurface nitrate is the dominant supply term, then the export should have
the same d15N as the sub-euphotic zone nitrate transferred to the sunlit surface
ocean. In contrast, within nutrient-rich regions of the surface ocean, where nitrate
consumption is persistently incomplete, the d15N of sinking N is lower than subeuphotic d15Nnitrate, and the d15N of the residual surface nitrate is higher than both
the sub-euphotic zone nitrate and the sinking N (Fig. 34.4, Southern Ocean)
(DiFiore et al., 2006; Lourey et al., 2003; Sigman et al., 1999b).
3.1.4. Deep nitrate
Despite dramatic variations of d15Nnitrate within the nutrient-bearing surface ocean,
nitrate assimilation has little effect on the global distribution of N isotopes. Complete
nitrate utilization in most of the surface ocean and virtually complete remineralization in the subsurface prevents the accumulation of isotopically distinct non-nitrate
pools. Dissolved organic nitrogen, the only other pool of appreciable size, has a very
similar isotopic composition to that of nitrate where it has been measured (Knapp
et al., 2005) and therefore can have little impact on the global distribution of 15N.
Nitrate concentrations in the deep sea are so high that the small amount of low-d15N
organic matter supplied from the nutrient-rich surface has only a minimal effect on
the d15N of deep ocean nitrate (Sigman et al., 2000) (compare Southern Ocean deep
water vs. Sargasso Sea, Fig. 34.4).
Instead, the modern deep sea is remarkably homogeneous with nitrate d15N of
5 þ/ 0.5%, so that it should represent a well-integrated measure of global input
and output rates, rather than internal cycling. Mean deep ocean d15Nnitrate therefore
has the potential to provide a monitor of whole-ocean inventory change.
3.2. Ocean d15Nnitrate through time
It seems safe to assume that N2 fixation, water column denitrification and sedimentary denitrification have been the dominant input/output terms of the marine N
budget throughout, at least, the past 2 Myr. With this assumption in place, one can
begin to consider how potential changes would have altered the d15N of marine
nitrate, both within a given region and at the scale of the whole ocean.
3.2.1. Local variations
Because the isotope fractionating processes are not evenly distributed throughout the
ocean, temporal variations at fixed points will produce predictable changes in the
d15N of local nitrate pools. As examples: (1) an increase in denitrification rates within
the water column on the Oman margin would increase the d15Nnitrate of the Arabian
Sea thermocline (Altabet et al., 1995; Schafer and Ittekkot, 1993), (2) an increase
in N2-fixation in surface waters of the eastern Mediterranean would cause a decrease
in d15Nnitrate in the eastern Mediterranean thermocline (Struck et al., 2001), and (3)
an increase in relative nitrate consumption in Antarctic surface waters would cause an
increase in local near-surface d15Nnitrate (Francois et al., 1997). However, it must be
kept in mind that these local signals are imprinted upon any variation in the d15N
baseline of mean ocean nitrate. A similar principle operates at finer scales, with
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changes in a N-cycle process in one region contaminating adjacent regions to various
degrees (e.g., Kienast et al., 2002; Martinez et al., 2006; Galbraith et al., 2008).
3.2.2. Steady state mean
At steady state, the flux-weighted mean d15N of fixed N outputs (dominated by
water column and sedimentary denitrification) must equal that of the inputs (dominated by marine N2 fixation). The d15N of nitrogen supplied by N2 fixation is tied
to the d15N of the atmosphere, which does not vary on a million year time scale. The
d15N of nitrogen removed by water column and sedimentary denitrification, on the
other hand, is tied to the mean d15N of oceanic nitrate, subject to the effective isotope
discriminations. That the mean ocean d15Nnitrate is higher than the d15N of atmospheric N2 is fundamentally due to isotope discrimination during denitrification.
At steady state, the mean ocean d15Nnitrate approaches the value at which total
denitrification (occurring in both the water column and sediment) removes nitrate
with the same d15N as N added to the ocean by N2 fixation, i.e. 1%. Given this
condition, the ratio of rates of sedimentary versus water column denitrification
provides the overarching control on the mean d15N of nitrate (Brandes and
Devol, 2002). Roughly speaking, since water column denitrification discriminates
against 15N-nitrate while sedimentary denitrification does not, a higher ratio of
water column to sedimentary denitrification will tend to yield a higher mean nitrate
d15N (Fig. 34.5). For example, if water column denitrification were homogeneously
distributed throughout the oceans, an expressed isotope effect of 25% for water
column denitrification would require that sedimentary denitrification be responsible
for roughly 80% of the modern loss of fixed nitrogen from the ocean in order to
maintain a mean ocean nitrate d15N of 5% (Brandes and Devol, 2002). The fact that
water column denitrification takes place in <0.2% of the ocean’s volume, with
relatively high degrees of nitrate loss (typically exceeding 40%), means that water
column denitrification causes less 15N enrichment than would arise if water column
denitrification were homogeneously distributed. The first effort to account for this
‘‘dilution effect’’ (Deutsch et al., 2004) indicates that the preindustrial balance would
require 30% of denitrification to take place in the water column (i.e., 50% more
than if the ocean were approximated as homogeneous). But such quantitative
estimates remain preliminary; there is more to be learned about the net isotopic
effect of both water column and sedimentary denitrification and their potential to
vary over time.
3.2.3. Application to the past
As mentioned above, a given sediment core will simultaneously provide information
on at least three separate quantities: changes in mean ocean nitrate d15N, changes in
regional 15N depletion or enrichment relative to the whole ocean, and changes in
the completeness of nitrate consumption in the local surface layer. Obviously, the
existence of a record of past variations in the mean ocean d15Nnitrate would allow
quantitative assessment of local and regional phenomena in other records; however,
the global background signal has yet to be explicitly resolved from among the array
of local changes. Sediment cores from the low and mid-latitudes contain the remains
of plankton grown on N supplied from the thermocline and not directly from the
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Figure 34.5 Mean ocean steady-state N isotope balance.Vertical arrows represent global input and
output fluxes, in which the flux is proportional to the length of the arrow and the isotopic composition is shown by the position on the horizontal axis (showing d15N in %). The yellow
downward-pointing arrow represents the N2 fixation flux, the white arrow represents the sedimentary denitrification flux, and the blue arrow represents the water column denitrification flux.
Sedimentary denitrification removes N of d15N close to that of mean ocean nitrate.Water column
denitrification removes nitrogen depleted in 15N by the expressed isotope effect, E (red arrow).
The green reservoir below the axis represents the d15N of mean ocean nitrate, which is determined
by the ratio of sedimentary to water column denitrification. At steady state, mass and isotopic balance requires that the sum of the denitrification fluxes equal the N2 fixation flux and that the denitrification fluxes be balanced about the N2 fixation ‘‘fulcrum.’’ Hence, a balanced isotope budget
with an expressed isotope effect of 25% for water column denitrification requires a much larger
flux of sedimentary denitrification in order to maintain a mean ocean d15N of 5%, close to the fulcrum (the d15N of N2 fixation, 1%), as shown by (A). Keeping the total N2 fixation flux equal
and increasing the relative proportion of water column denitrification to equal that of sedimentary denitrification pushes the mean ocean reservoir to a higher value (B). If the expressed isotope
effect of water column denitrification is less than 25%, the mean ocean nitrate would be less sensitive to changes in the relative proportion of water column denitrification, but deviations from
5% would still be possible (C, D).
deep sea; given the regional heterogeneity of thermocline d15Nnitrate (which exhibits
a 15% range in the modern ocean), it is difficult to extract the global mean
background variability from a foreground of changes in surface and thermocline
nitrate d15N. Although polar ocean regions sample deep nitrate more directly, the
possibility that the degree of nitrate consumption has varied means that records from
these regions cannot be interpreted solely in terms of the deep nitrate supply. Before
exploring this further, however, we must discuss the development of sedimentary
records of the N isotopes.
3.3. Sedimentary nitrogen isotope records
The most common paleoceanographic proxy used to reconstruct past changes in the
oceanic N cycle is the d15N of total combustible, or ‘‘bulk,’’ sedimentary nitrogen
(hereafter d15Nbulk). Such measurements provide a record, with a variable degree of
accuracy, of the N sinking out of the surface ocean at times past (Altabet and Francois,
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1994). Below, we review the robustness of the connection between the organic N
sinking out of the surface ocean and that preserved in paleoceanographic records as
bulk sedimentary N. We also briefly discuss the ongoing development of alternatives
to the bulk measurement.
3.3.1. The d15N of sinking nitrogen relative to local nitrate
In low latitude regions where nitrate is nearly completely consumed, sedimentary
d15N records are typically thought to reconstruct the d15N of the sub-euphotic zone
nitrate that is supplied to the euphotic zone by upwelling and/or mixing. But, in
some cases, the d15N of the sinking flux may not match the d15N of the nitrate
supplied to the euphotic zone. For instance, in some regions, diazotrophic biomass
may account for a significant fraction of the sinking N, such that the d15N of N
sinking out of the surface ocean may be significantly lower than the subsurface
d15Nnitrate (Brandes et al., 1998; Horikawa et al., 2006; Karl et al., 1997; Pantoja et al.,
2002). As discussed in Section 3.1.3, in nutrient-rich regions such as the Antarctic
and Subarctic Pacific, the gross nitrate supply to the surface is only partially consumed before being mixed or subducted back into the ocean interior or transported
laterally out of the system. In this case, the d15N of sinking organic matter will be
lower than the d15N of the nitrate initially supplied to that region of the surface
ocean (Altabet and Francois, 1994, 2001; Sigman et al., 1999b; Wu et al., 1999). It
has been suggested that the progressive consumption of nitrate can also be gleaned
from sedimentary transects underlying coastal upwelling cells (Holmes et al., 1996),
though delicate gradients such as these would be prone to obliteration by lateral
sediment transport on the slope (Mollenhauer et al., 2003) and should be interpreted
with caution.
Beyond these upper ocean processes, poorly understood phenomena alter the
d15N of sinking nitrogen during transit through the water column and burial. In a
variety of oceanographic settings, sinking particulate d15N has been observed to
decrease with depth in the water column (by up to 2% over several kilometers of
depth) (Altabet and Francois, 1994; Lourey et al., 2003; Nakanishi and Minagawa,
2003). Explanations for this range from the addition of low-d15N bacterial biomass,
to the contamination of shallower traps by zooplankton ‘‘swimmers’’. More recently,
other deviations from expectations have been observed. The sinking nitrogen d15N
in the Southern Ocean is highest during the wintertime low flux period and then
decreases into the higher flux, spring bloom period as [NO3 ] declines (Altabet and
Francois, 2001; Lourey et al., 2003). This sense of change is the opposite of what is
expected from simple isotopic fractionation during algal assimilation of surface
nitrate. It may reflect domination of the wintertime sinking flux by very slowly
sinking material that has been enriched in 15N by progressive degradation, analogous
to observations of high d15N in deep suspended particles (Saino and Hattori, 1987).
3.3.2. The d15N of sedimentary nitrogen
Despite these observations, excellent correlation is found between the d15N of bulk
surface sedimentary N and that of the local sub-euphotic zone nitrate at sites where
export production is high and/or organic matter preservation is good (many of
which are from continental margins, Fig. 34.6) (Thunell et al., 2004). In these
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Figure 34.6 Fidelity of sedimentary d15N as a recorder of d15Nnitrate at coastal margins. This plot
shows the d15Nnitrate of water at the base of the euphotic zone versus the d15N of surface sediment
at or near the same locations. Despite the complexity of nitrogen cycling within the water column,
the sediment records the d15N of thermocline nitrate with a high degree of accuracy at all of these
diverse locations. Data from (Altabet et al., 1999b; Emmer and Thunell, 2000; Kienast et al., 2002;
Liu, 1979; Pride et al., 1999; Thunell et al., 2004), figure modified after Thunell et al., (2004).
settings, alteration during sinking and incorporation in the sediments does not
appear to confound bulk sedimentary N as a recorder of the d15N of the N sinking
out of the surface ocean. It also attests to the generally tight connection between the
d15N of the nitrate supply to the surface and that of the annually integrated sinking
flux under conditions of essentially complete nitrate consumption in surface waters
(Altabet et al., 1999b).
Measurements of d15Nbulk from slowly accumulating sediments at the deep sea
floor have uncovered spatial patterns that qualitatively mirror expectations for the
spatial variation in the d15N of sinking N (Altabet and Francois, 1994; Farrell et al.,
1995; Holmes et al., 1996). However, in contrast to the ocean margin settings, a
significant 15N enrichment (on the order of 1–5%) is observed in bulk sedimentary N
at core tops relative to sinking particles (Altabet and Francois, 1994; Gaye-Haake
et al., 2005). This d15N increase is apparently due to preferential remineralization of
14N during the oxidation of organic matter in sediments, and the isotopic effect of this
process appears to be greatest in organic-poor deep-sea sediment, where accumulation rates are low (Prokopenko et al., 2006). As one might expect, there is some
evidence for a continuous range in the d15N difference between sinking and sedimentary N, from no difference in the high-preservation margin environments
described above to 5% in the most organic-poor, slowly accumulating sediments
of the open ocean (Nakanishi and Minagawa, 2003; Pichevin et al., 2007). As a result,
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there is concern that pronounced changes over time in sedimentation rate, or in the
conditions at the sediment-water interface, could produce a time-varying degree of
diagenetic enrichment of sediment d15N (Sachs and Repeta, 1999).
Sedimentological processes introduce additional uncertainties that deserve mention. Two sedimentary d15N records from adjacent sites on the northern margin of
the South China Sea show quite different features over the past 40 kyr despite being
separated by less than 15 km (Higginson et al., 2003; Kienast, 2005). This is apparently
due to sediment transport that redistributes the fine fraction of sediment, including
the organic matter in which the N resides, over large distances (Freudenthal et al.,
2001; Kienast, 2005). Other complicating factors can include the presence of terrestrial organic matter, which may have a distinct N isotopic composition (McKay et al.,
2004; Peters et al., 1978), and inorganic N, which can be a significant component in
the lattices of some clays, particularly illite (Schubert and Calvert, 2001).
Because of concerns about seafloor processes affecting the d15N of bulk sedimentary N, the isotopic composition of targeted sedimentary N fractions is being
pursued. Most work of this type has focused on the N bound within the silica
frustules of diatoms, which should record past surface ocean processes while being
protected from the isotopic effects of bacterial degradation during sinking and
incorporation in the sedimentary record (Crosta and Shemesh, 2002; Robinson
et al., 2004; Shemesh et al., 1993; Sigman et al., 1999a). The d15N of organic matter
in foraminiferal tests has been explored following the same rationale (Altabet
and Curry, 1989). Chlorophyll degradation products have also been isolated from
sediments for N isotopic analysis (Ohkouchi et al., 2006; Sachs and Repeta, 1999).
The ongoing research into such specific N fractions is critical to address the
fundamental issue of isotopic alteration of exported organic N in the water column
and sediments. This work also raises a host of new questions. Consider the case of
diatom frustule-bound N (Fig. 34.7). The organic N of diatoms exported from the
surface layer could differ from the integrated sinking N, for instance, if the isotope
effect of nitrate assimilation by diatoms differs from that of the entire phytoplankton
population. Moreover, the d15N of N trapped and preserved within the diatom
frustule (here d15Nfrustule) appears to be quite different from the bulk biomass of the
diatom that produced it. Thus, d15Nfrustule need not be a direct reflection of
integrated sinking flux d15N, and the relationship between the two could change
over time (Galbraith, 2006). Nevertheless, the concerns regarding the interpretation
of bulk d15N demand that approaches be developed for the analysis of specific N
fractions, at least for open ocean records. The combination of bulk d15N with
specific N fractions from the same sediment samples may yield important complements in the effort to reconstruct changes in the N cycle.
4. Observations and Interpretations
Although questions remain regarding the fidelity of sedimentary N isotopes
as recorders of oceanic N dynamics, significant progress has been made by comparing multiple records from different sedimentary regimes and by using additional
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δ15Nexport
Δ(d-e)
N export
Δ(f-d)
Δ(diagenesis)
?
+ 0-5‰
δ15Nfrustule
δ15Nbulk
Figure 34.7 The d15N of integrated N export (represented by the downward pointing arrow,
‘‘d15Nexport’’) versus d15Nbulk and d15Nfrustule, two sedimentary measurements intended to reconstruct the d15N of N export. During sinking and burial in slowly accumulating deep-sea sediments, organic N may undergo 15N enrichment or depletion owing to diagenesis. The amplitude
of diagenetic alteration is environmentally variable, being virtually absent in organic-rich ocean
margin environments but as high as 5% in the less productive open ocean. The d15N of N bound
within diatom frustules (d15Nfrustule) does not appear to be affected by diagenesis. However,
measurements to date indicate that it is different than ‘‘d15Nexport’’. This difference may be due to
(1) a d15N difference between diatom biomass sinking out of the euphotic zone and the bulk
sinking N (‘‘D(d-e)’’), and/or (2) a d15N difference between the organic N incorporated in diatom
frustules and the bulk biomass N of the diatom (‘‘D(f-d)’’). The amplitude of either of these d15N
differences may be sensitive to changes in surface ocean conditions (e.g., because of changes
in phytoplankton assemblage or in the growth conditions of diatoms). Thus, given current
information, neither d15Nbulk nor d15Nfrustule can be assumed equal to d15Nexport.
proxies to better constrain other environmental parameters. Indeed, many of the
records generated to date show coherent variations on different spatial and temporal scales, indicating that they are undoubtedly recording meaningful signals
with strong links to global climate. Here we review some of the observations to
date and their current interpretations, in relation to the conceptual mechanisms
outlined in Section 2.
4.1. Glacial–interglacial inventory changes
The most robust finding to arise from sedimentary N isotope records is of
climatically-linked changes in d15Nnitrate (Fig. 34.8) near each of the major thermocline oxygen minimum zones of the global ocean: the Arabian Sea and the eastern
tropical North and South Pacific (Altabet et al., 1995; Ganeshram et al., 1995, 2000;
Pride et al., 1999). Although it remains possible that these signals represent changes
in the degree to which 15N-enriched nitrate from denitrification zones invaded the
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Figure 34.8 Plots of d15Nbulk versus sediment age over the past 50,000 years for cores from the
Mexican margin (Ganeshram et al.,1995), the Somali margin of the Arabian Sea (Ivanochko et al.,
2005) and the South China Sea (Kienast, 2000). The Last Glacial Maximum (LGM) and Younger
Dryas (YD) cold intervals are indicated by the blue shading. The d15N offsets among the records
over the Holocene (the last 10 kyr) are consistent with the effect of modern water column denitrification, which causes high thermocline d15Nnitrate along the Mexican margin and in the Arabian Sea. During the last glacial maximum, this range collapsed to roughly half that of the
present day (2%).This appears to be due to a reduction of water column denitrification, reducing
the preferential loss of 14N from the oceans.The remarkable stability of the South China Sea d15N
record in the face of this glacial reduction in water column denitrification supports compensatory
global-scale reductions in both sedimentary denitrification and N2 fixation.
neighboring water masses (Kienast et al., 2002), we follow here the common
assumption that the records are dominated by past changes in the extent and/or
activity of the thermocline denitrification zones. In addition, both the west African
margin and the subarctic Pacific, which have intense thermocline O2 minima but
lack denitrification today, yield similar sediment d15N records, suggestive of water
column denitrification during past periods (Galbraith et al., 2004; Lavik, 2002;
Pichevin et al., 2005). We suggest three factors supporting the assumption that
these records document variable denitrification in the thermocline. First, the modern thermocline denitrification zones are all associated with highly productive
margin environments. As described in Section 3.3.2, the sediments from these
environments have been shown to accurately record the d15N of subeuphotic
zone nitrate (Altabet et al., 1999b; Thunell et al., 2004), making many of them
good monitors of nearby denitrification in the thermocline. Second, the isotope
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effect of water column denitrification is large, so that the isotopic signal of denitrification readily overcomes the noise of other processes (see ETNP, Fig. 34.4). Third,
a number of other sediment properties (e.g., presence of laminations, redox-sensitive
metal contents) corroborate regional variations in water column oxygenation that
parallel the inferred denitrification changes.
All of the denitrification records show a general pattern of low d15N during
glacial periods and high d15N during warm interglacial periods (Fig. 34.9). These
changes are accompanied by parallel evidence of water column suboxia from trace
metal concentrations (Hendy and Pedersen, 2005; Ivanochko and Pedersen, 2004;
Kienast et al., 2002; McKay et al., 2005; Nameroff et al., 2004), benthic foraminiferal
assemblages (Ohkushi et al., 2003) and sediment laminations (Ganeshram et al., 2000;
Suthhof et al., 2001; Thunell and Kepple, 2004; van Geen et al., 2003), confirming
that thermocline suboxia was expanded during periods with high d15N. The inverse
correlation between denitrification in the ETNP and global ice volume has been
shown to have reigned throughout, at least, the past 1.8 My (Liu et al., 2005). The
deglacial increase in thermocline denitrification does not appear to have been
contemporaneous everywhere; in particular, records from Peru (Higginson and
Altabet, 2004), Chile (De Pol-Holz et al., 2006; Robinson et al., 2007; Martinez
et al., 2006) and within the core of the ETNP shadow zone (Hendy and Pedersen,
2006) show d15N rises prior to 16 kybp, while most well-dated records from near
the ETNP and Arabian Sea increase rapidly after 15 kybp. Nonetheless, on a glacialinterglacial timeframe, there appear to be grossly coherent changes in integrated
denitrification rates in the global thermocline, with local variability modulated by
regional-scale fluctuations of circulation and productivity.
The globally recognized, millennial-timescale climate oscillations that punctuate
the broader glacial cycles (Dansgaard et al., 1993) also exhibit significant changes in
denitrification. For example, during the Younger Dryas, a brief return to ice-age-like
conditions that interrupted the last deglaciation between 12.9 and 11.5 kybp, there
was less denitrification in the ETNP (Hendy et al., 2004; Ivanochko and Pedersen,
2004) and Arabian Sea (Altabet et al., 2002; Higginson et al., 2004; Ivanochko et al.,
2005; Suthhof et al., 2001). Increased oxygen concentrations at intermediate depth
are corroborated by decreased sediment lamination, lower trace metal concentrations, and decreased age of intermediate waters (Ahagon et al., 2003; Behl and
Kennett, 1996; Ivanochko and Pedersen, 2004; Keigwin et al., 1992; Pride et al.,
1999; van Geen et al., 2003). Earlier, between 25 and 60 ky ago, the DaansgaardOeschger climate oscillations seem to have involved similar changes in denitrification
intensity within both the Arabian Sea (Altabet et al., 2002; Ivanochko et al., 2005;
Suthhof et al., 2001) and the North Pacific (Emmer and Thunell, 2000; Hendy et al.,
2004). Although determining precise chronologies for sediments in this time range is
difficult, the oscillations appear to be in phase with climate records from the North
Atlantic, with warm conditions in the North Atlantic co-occurring with enhanced
denitrification. In contrast, denitrification records in the Southern Hemisphere
appear to be roughly antiphased to those of the north, so that denitrification is instead
in phase with Antarctic warm events (Martinez et al., 2006; De Pol-Holz et al., 2006;
Robinson et al., 2007).
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Antarctic Zone
opal preserved flux (Th-norm, g cm–2kyr –1)
0.0 1.0 2.0 3.0
0
diatom-bound
bulk sediment
50
depth (cm)
100
150
LGM
200
250
300
350
0
2
4
6
8
δ15N (‰ vs. air)
2.5 3.0 3.5 4.0
plank. δ18O (‰ vs. PDB)
opal preserved flux (Th-norm, g cm–2kyr –1)
Subantarctic Zone
0.0 1.0 2.0 3.0
0
depth (cm)
100
200
300
LGM
400
500
600
0
2
4
6
8
δ15N (‰ vs. air)
2.0 2.5 3.0 3.5
plank. δ18O (‰ vs. PDB)
Figure 34.9 Paleoceanographic records from the Antarctic (MD84-552) and Subantarctic
(MD84-527) zones of the Southern Ocean. Foraminiferal d18O is shown as a proxy for global ice
volume, to indicate the Last Glacial Maximum, which is shaded gray. The 230Th-corrected opal
flux is shown in the centre of each panel as a proxy for diatom export production, and both
d15Nbulk and d15Nfrustule are shown on the left (open and filled circles, respectively). During the
LGM, diatom export was lower in the Antarctic Zone, and virtually unchanged in this (Indian)
sector of the Subantarctic Zone.The d15Nbulk records appear to be affected by diagenetic alteration
in this well-oxidized, low-accumulation-rate environment. However, the d15Nfrustule shows that
the d15N of diatoms was higher in both zones during the LGM. The simplest way to reconcile
these data is to invoke, during the LGM, decreased upward macronutrient supply throughout the
surface of the Antarctic. After Robinson et al., 2004, 2005.
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The dependence of water column suboxia on sinking organic matter makes
variable export productivity a clear contender for control of denitrification (Altabet
et al., 1995; Ganeshram et al., 1995), and there is evidence for increased export
production during warm periods in oxygen-sensitive regions (Altabet et al., 1999a;
Crusius et al., 2004; Ganeshram et al., 2000; Hendy et al., 2004; Ortiz et al., 2004).
Such changes have been most commonly ascribed to changes in local wind-driven
upwelling, although changes in thermocline preformed nutrient concentrations
(Robinson et al., 2005) and the global nutrient distribution (Boyle, 1988;
Schmittner, 2005, 2007) may also play an important role, which we return to in
section 4.3 below. An enhanced supply of oxygen to poorly ventilated subsurface
regions, through a combination of lower water temperatures (higher O2 solubility)
and accelerated circulation of thermocline water masses, would also have contributed
to reducing glacial denitrification rates (Fig. 34.9, (Galbraith et al., 2004; Kennett and
Ingram, 1995; Meissner et al., 2005)). A global coupled ocean-ecosystem model has
shown that both reduced oxygen consumption and increased intermediate-depth
oxygen supply can arise from a shutdown of NADW, as might be induced by a
massive iceberg discharge to the North Atlantic region (Schmittner et al., 2007). This
has the potential to explain the Dansgaard-Oeschger and Younger Dryas observations from the Northern Hemisphere, and hints that the antiphased Southern Hemisphere denitrification records might reflect correspondingly pronounced changes in
the southern oceanic circulation.
Despite the large reorganization of Earth’s climate system and apparent changes
in thermocline denitrification, the data in hand do not suggest a clear change in
mean ocean d15Nnitrate between the LGM and present day (Altabet and Curry, 1989;
Kienast, 2000). This constraint is based on the observation of little glacial-interglacial
change (<0.5%) in many cores, including in the North Atlantic (Huon et al., 2002),
Southern Indian Ocean (Francois et al., 1997), Sulu Sea (Horikawa et al., 2006) and
South China Sea (Kienast, 2000; Tamburini et al., 2003). This is in stark contrast to
marine oxygen isotopes, for example, which are overwhelmed by a global glacialinterglacial signal (Lisiecki and Raymo, 2005). The South China Sea records appear
to be compromised to some degree by sedimentary processes (Kienast, 2005), which
may mask a glacial-interglacial change to some degree. More importantly, all lowlatitude records are highly sensitive to the opposing, dramatic isotopic signals of
denitrification and N2 fixation that permeate the thermocline (Galbraith et al.,
2004); lack of variability at any given site could be due to fortuitous location at an
isotopic balance point between regions of denitrification and N2 fixation. Furthermore, as noted in Section 3.3.2, some of these deep-sea records are likely affected by
diagenesis, which casts doubt on their fidelity as recorders of the glacial-interglacial
change. Nonetheless, lack of an obvious global change between the LGM and
Holocene is an important constraint, given that the short residence time of marine
N (3 kyr, (Codispoti et al., 2001)) requires that steady state was approached during
the LGM and has been approached again since the deglaciation. Under the reigning
N isotope paradigm, the global d15Nnitrate during the LGM could only have been
similar to that of today if the ratio of water column to sedimentary denitrification
during the LGM was also similar to the present day (see Section 3.2.2, Fig. 34.5).
Given the strong evidence for lower water column denitrification rates during the
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glacial, this implies that the global rate of sedimentary denitrification was significantly decreased as well, in agreement with the expected effect of lower glacial sea
level (see Section 2.1.2). As a result, the total glacial denitrification rate must have
been substantially lower than today.
Given the constraint that the LGM N budget was near steady state, we can infer
that the global rate of N2 fixation was lower, by an amount roughly equal to the
aggregate decrease in denitrification. This corresponding decrease in N2 fixation
does not preclude changes in the N inventory, however; a small sustained surplus in
the N budget, possibly abetted by an increased P inventory or trace-metal fertilization of diazotrophs (see Section 2.1.1), could have allowed for a greater N inventory
during the glacial period, subject to the strength of the N2 fixation self-limiting
feedback (see Section 2.2.1). All things considered, it would seem likely that the
glacial N inventory was somewhat higher, though a quantitative estimate is not
currently possible. A tentative upper bound, provided by a numerical analysis of
observed deglacial d15N patterns, constrains the glacial inventory to no more than
30% larger than the present day inventory, despite the possibility of much larger
shifts in fluxes (Deutsch et al., 2004). More precise quantification of the inventory
change will require a more global distribution of d15N records, better understanding
of the denitrification isotope effects, and numerical modeling.
It is possible that local manifestations of the deglacial acceleration of N2 fixation
left traces in sediments, and records from the Cariaco Basin provide a tantalizing
candidate for this (Haug et al., 1998). However, N2 fixation within the Cariaco
Basin may respond to local, as well as global changes in denitrification (Haug et al.,
1998; Meckler et al., 2007), and no published open ocean records show a sustained
Holocene decrease in d15N below LGM values. The lack of clear N2 fixation records
may result from a close physical proximity of N2 fixers to denitrification zones
(Brandes et al., 1998; Sigman et al., 2005; Westberry et al., 2005; Deutsch et al.,
2007), such that the smaller isotopic signal of N2 fixation is overwhelmed by the
amplitude of the denitrification signal. Otherwise, it may be the lack of good deepsea records from oligotrophic gyres that is hindering our view of N2 fixation. The
resolution of this issue is bound up in the current debate over where the bulk of
modern N2 fixation occurs (Capone et al., 2005). We note that although clear
records of nitrogen fixation from the open ocean are elusive, dramatic reductions
of d15N within the organic-rich sapropels that punctuate Quaternary sediments of
the Mediterranean (Calvert et al., 1992) have been shown to reflect large increases in
nitrogen fixation (Struck et al., 2001). Similar d15N reductions within ancient
organic-rich sediments have been interpreted as reflecting analogous periods of
intense N2 fixation during the Cretaceous (Kuypers et al., 2004; Rau et al., 1987).
Intriguingly, despite the lack of a global LGM-Holocene change, a deglacial
maximum in d15N is observed in the majority of sediment records, followed by a
d15N decrease during the Holocene (Bertrand et al., 2000). Given the arguments for
large variations in water column denitrification, the most straightforward explanation
for the maximum invokes a peak in water-column denitrification rates, followed by a
decrease during the Holocene. This sequence could conceivably have been driven by
either an environmental change or progressive strengthening of the negative feedback
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on water column denitrification through local nitrate limitation (Section 2.2.3;
Deutsch et al., 2004; Galbraith et al., 2004). The deglacial d15N maximum is such a
pervasive signal in sediment records, however, that it may reflect a transient maximum
in the mean ocean d15Nnitrate itself, caused by a maximum in the global ratio of water
column to sedimentary denitrification (Deutsch et al., 2004). Indeed, this could have
arisen simply due to a temporal offset between early and rapid increases in water
column denitrification (17–14 kybp) vs. the more gradual rise of sea level and flooding
of the continental shelves (18–10 kyr). Nevertheless, the largest d15N maxima occur in
sediments close to the denitrification zones, suggesting that the absolute rate of
thermocline denitrification also peaked during the deglaciation and has decreased to
some degree since then.
The pronounced changes in thermocline oxygenation that have accompanied
glacial-interglacial climate changes also imply large changes in the marine flux of
N2O to the atmosphere. The production of N2O in the ocean is greatly increased
under low oxygen concentrations, and oxygen-poor waters broadly associated with
denitrification dominate the marine source (Suntharalingam et al., 2000). Given that
the modern marine N2O flux is approximately one third of the global source (Hirsch
et al., 2006), any large change in water column denitrification related to a regional
change in subsurface oxygenation should have a significant impact on atmospheric
N2O concentrations. Rapid changes of atmospheric N2O concentrations recorded
in ice cores, including a sharp rise 14.6 ky ago as well as millennial-scale fluctuations
during the last ice age (Fluckiger et al., 1999, 2004) may have been largely caused by
the same changes of thermocline oxygenation that resulted in variable denitrification
rates (Suthhof et al., 2001).
In summary, the globally aggregated denitrification rate was generally lower
during glacial periods, oscillated on millennial timescales in concert with climate,
and accelerated rapidly during deglaciations. The emergent conclusion is that the
marine nitrogen budget is highly dynamic, responding to changes in the physical
environment, yet its input and output terms are coupled, leading to a nearly balanced
budget on timescales of centuries to millennia. Whether this coupling allows for
significant changes in the ocean N reservoir is a matter of debate. As described
earlier, the apparent lack of a glacial-interglacial global d15N change, and the
characteristics of certain deglacial signals in the N isotope records, suggest to us
that the ice age ocean N reservoir was no more than 30% greater than that of the
Holocene (Deutsch et al., 2004).
4.2. Glacial–interglacial changes in nutrient-rich regions
Nutrient-rich regions of the surface ocean have undergone significant changes in
export production over glacial-interglacial cycles. In the Antarctic Zone of the
Southern Ocean, numerous studies have converged to show that export production
was lower during the ice ages than during interglacials (Anderson et al., 1998;
Francois et al., 1997; Kohfeld et al., 2005). In contrast, studies of the Subantarctic
Zone generally indicate higher export production during ice ages, the inverse of the
Antarctic climate cycle in productivity (Chase et al., 2003; Frank et al., 2000; Kumar
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et al., 1995; Mortlock et al., 1991). In the Equatorial Pacific, reconstructions of
glacial-interglacial export productivity are conflicting, and the picture remains
equivocal (Farrell et al., 1995; Loubere, 1999; Loubere et al., 2004; Marcantonio
et al., 2001; Paytan and Kastner, 1996; Pedersen, 1983), while in the subarctic
Pacific, export productivity seems to have been generally lower during glacial
periods (Brunelle et al., 2007; Jaccard et al., 2005; Kienast et al., 2004; Narita et al.,
2002), similar to observations from the Antarctic. The observation of variable export
productivity in these nutrient-rich regions prompted sedimentary N isotope studies
to explore the underlying physical mechanisms, assuming that d15N variability there
would be dominated by changes in the degree of relative nitrate consumption.
In the glacial Antarctic, evidence for reduced export production is accompanied
by higher sedimentary d15Nbulk (Fig. 34.9) (Francois et al., 1997), implying a greater
degree of relative nitrate consumption despite the diminished export production.
Although d15Nfrustule records suggest that the glacial/interglacial change in environmental conditions was less homogeneous across the Antarctic zone than indicated by
d15Nbulk (Robinson et al., 2004; Robinson and Sigman, 2008), they remain consistent with equal or greater relative nitrate utilization during the glacial. This requires
that the gross rate of nitrate supply to the Antarctic euphotic zone was reduced
during glacial times. Physically, this change could be explained by reduced exchange
of water between surface and deep waters, as would result from an increase of
density gradients or a change in the wind-driven circulation (see Section 2.3).
Similarly, increased d15Nbulk and d15Nfrustule in the subarctic Pacific co-occur with
decreased export production during glacial periods, mirroring the Antarctic trend
(Brunelle et al., 2007; Galbraith et al., 2008), consistent with theories of glacial water
column stratification in the polar oceans (Haug et al., 1999; Jaccard et al., 2005;
Sigman et al., 2004). Enhanced aeolian Fe supply during the glacial may have
contributed to nitrate drawdown by supplementing the Fe from deep waters
(Archer and Johnson, 2000; Lefevre and Watson, 1999; Parekh et al., 2004), but
this cannot be resolved from the current data.
Although d15Nbulk measurements from the Subantarctic show few systematic
glacial/interglacial patterns of change (Francois et al., 1997), d15Nfrustule records from
the Subantarctic reveal a coherent pattern of elevated d15N during glacial times (Fig.
34.9) (Crosta et al., 2005; Robinson et al., 2005). The combination of increased opal
fluxes with elevated d15N is consistent with higher nutrient consumption as
expected under enhanced Fe supply. Given the greater proximity of the Subantarctic
Zone to potential dust sources in Patagonia, South Africa, and Australia, and the
zonality of the winds, the d15N may indicate a greater susceptibility to iron fertilization during ice ages than the higher latitudes of the Southern Ocean (Robinson
et al., 2005).
In the equatorial Pacific, d15Nbulk is low during glacial episodes and elevated
during interglacials (Farrell et al., 1995). While this may indicate lower nitrate
consumption during the ice age, there are two important caveats to consider.
First, the correlation of low d15Nbulk with periods of rapid sediment accumulation
is reminiscent of the Subantarctic records, in which d15Nfrustule data have suggested a
diagenetic origin for the d15Nbulk changes (Robinson et al., 2005). Second, the
equatorial Pacific exchanges water with the ETNP shadow zone (Higginson et al.,
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2003; Yoshikawa et al., 2005), so the apparent glacial decrease in denitrification
within the ETNP may have lowered the regional d15Nnitrate enough to impact the
d15N of glacial-age sediment in the equatorial Pacific. Thus, the nutrient status of the
ice-age equatorial Pacific remains enigmatic.
4.3. Connections
Although we have presented inventory changes and nutrient utilization as separate
processes, they are undoubtedly linked. Perhaps most interesting is the potential role
of high latitude circulation and nutrient utilization on the nutrient content of the
thermocline, a key parameter for setting the productivity of low latitude surface
waters (Sarmiento and Orr, 1991; Sarmiento et al., 2004; Schmittner, 2005).
The nutrient cycling scenarios presented above show the potential for Southern
Ocean surface waters to become relatively nutrient-poor under alternate climate
states. If such nutrient depletion did indeed occur in the glacial ocean, the net result
would likely have shifted nutrients from the surface and thermocline to the deep sea
(Boyle, 1988; Keir, 1988). Glacial decreases in the formation rate of North Atlantic
Deep Water (NADW), intimately linked to North Atlantic surface conditions,
would also have decreased the rate at which nutrients were returned from the
deep ocean to the surface (Schmittner, 2005). Both of these mechanisms fit with
extensive evidence from foraminiferal d13C and Cd/Ca for thermocline nutrient
depletion during the last ice age, in the Atlantic (Marchitto et al., 1998), the North
Indian Ocean (Boyle, 1988, 1992; Kallel et al., 1988), and the Pacific (Herguera
et al., 1991; Keigwin, 1998; Matsumoto et al., 2002). The loss of nutrients from the
sunlit surface ocean would have driven a decrease in low latitude export production,
for which there is considerable evidence (Ganeshram and Pedersen, 1998; Loubere
et al., 2004; Marcantonio et al., 2001; Ortiz et al., 2004; Ivanochko et al., 2005). This,
in turn, would have reduced oxygen consumption in the thermocline at the same
time as the subsurface oxygen supply was increased by physical means (Fig. 34.10).
Furthermore, global variations in the nutrient richness of the thermocline would
have caused sedimentary denitrification to vary in parallel with water–column
denitrification, maintaining a stable global mean d15Nnitrate as implied by the observations. Although this general mechanism remains speculative, it offers an intriguing
paradigm that is consistent with the evidence gathered to date.
Reflexively, it seems likely that changes in the subsurface nitrate d15N, driven by
low latitude processes, will become more apparent in high latitude records. For
instance, recent work has shown that sedimentary d15N records in the subarctic Pacific
represent the sum of broad, regional changes caused by denitrification in the tropical
Pacific, and local nutrient utilization within the subarctic itself (Galbraith et al., 2008).
The Antarctic, in contrast, is more intimately connected to the deep ocean and thus
may be most sensitive to d15N changes of deep nitrate. N isotopes aside, past changes
in the oceanic N reservoir could have affected relative nitrate consumption in
nutrient-rich regions of the ocean, as the gross nitrate supply to the surface would
have changed. Given that the low and high latitude ocean regions can affect one
another in these and other ways, future research will likely reveal complex relationships between the input/output budget for oceanic fixed N and its internal cycling.
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High latitude
Subsurface
O2 minimum
Warmer
Nutrient-rich surface
Widespread N2-fixation
Low [O2]
Excess P
Low
N:P
Slow circulation
Abundant denitrification
High nutrient : O2
Colder
Nutrient-poor surface
Little N2-fixation
Complete P
utilization
High
N:P
High [O2]
Rapid circulation
Little denitrification
Low nutrient : O2
Figure 34.10 Modulation of marine N budget by high latitude climate change.The colored layer
indicates the ventilated thermocline, in which darker blue represents greater phosphate concentrations. Downward-pointing green arrows represent export production, which consumes subsurface oxygen, while the upward pointing blue arrows represent the upwelling of nutrients. In
the upper panel, low relative nutrient utilization in high latitude source water regions produces a
nutrient-rich thermocline, fuelling abundant low-latitude export production. Meanwhile, warm
conditions tend to produce low initial O2 concentrations and the waters may circulate more slowly.
The high nutrient:O2 waters maintain abundant suboxia, with widespread denitrification and,
consequently, low N:P in thermocline water (i.e. waters with excess phosphate), which is subsequently upwelled to the surface. Algal growth in this upwelled water eventually yields nitratefree, phosphate-bearing water, which drives N2 -fixation. In contrast, the lower panel shows high
relative nutrient utilization in the source regions, producing a nutrient-poor thermocline that
limits export production at low-latitudes. In addition, cool conditions are likely to produce high
initial O2 and more rapid thermocline circulation. As a result, thermocline suboxia and denitrification are reduced. With low nutrient:O2 conditions inhibiting denitrification, the N:P ratio
in upwelling water will be high, encouraging complete P utilization by non-N2 fixing algae,
discouraging N2 fixation.
5. Perspectives on Progress
5.1. Looking backward
It may seem as though the current view of the past marine N cycle is highly
uncertain, but this should be viewed in the context of the last several decades.
The first proposition that climate change can influence the marine N cycle was
made by McElroy (1983), who pointed out that the short residence time of fixed
marine N (then believed to be 10 ky) made it a good candidate for change over
glacial-interglacial cycles. The marine nitrogen budget estimates of the day appeared to
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show a large deficit, so McElroy suggested that the present-day situation was part of a
grand 100 ky cycle: during ice age inception, advancing glaciers would have bulldozed terrestrial N into the ocean, while the erosion of continental shelves exposed by
sea level drop would have flushed N-rich sediments into the deep sea, to create a glacial
ocean rich with nitrogen in which phosphate was the limiting nutrient everywhere.
When shelf sediments were worn down and the ice sheets reached their maximum
extents, the supply of terrestrial N would have returned to something like its low,
modern rate. Denitrification would have eaten away at the accumulated surplus of N,
bringing us eventually to our modern-day position of a nitrogen-limited ocean with a
large N deficit. This hypothesis, which was novel and fully defensible at the time, was
predicated on the view of the marine N cycle as unresponsive, in that the marine fluxes
of fixation and denitrification were assumed to be ineffective at compensating for
changes in the external supply of N.
Over the past twenty years, modern process studies have taught us that the input
and output fluxes of the marine N cycle are dominated by marine microbes, not by
terrestrial inputs or sediment burial. Meanwhile, sedimentary N isotope studies have
shown that extensive changes in the d15N of thermocline nitrate have occurred, with
coherent connections to climate dynamics. These observations indicate that the
marine nitrogen cycle responds dramatically to changes in the physical environment
on timescales shorter than envisioned by McElroy, with large changes in thermocline
denitrification occurring within centuries. Moreveover, the realization that the residence time of marine nitrogen is only a few thousand years, coupled with the stability
of global mean nitrate d15N suggested by sediment records, requires that steady state is
roughly achieved on a thousand year time scale and not over tens of thousands of years.
These observations are consistent with a reactive, internally stabilized nitrogen cycle,
unlike the unresponsive, terrestrially controlled cycle described by McElroy.
Although quantitative interpretation of the paleoceanographic data has just
begun, the insights gained thus far suggest a tightly coupled system that reacts
dramatically to external forcings but is resistant to extreme inventory variations.
This, in turn, requires the existence of effective stabilizing feedbacks within the
marine N cycle. Thus, the remaining questions do not concern whether or not
these feedbacks exist, but rather how strong they are, how tightly they constrain the
system, and by precisely what mechanisms they operate.
5.2. Looking forward
Valuable information will continue to arise from measurements of the isotopic
composition of bulk sediments, especially in margin environments. Expanding the
network of high-resolution d15Nbulk records is crucial to map out patterns of
changing d15Nnitrate in space and in time, without which mechanistic hypotheses
cannot be tested. However, it is critical that we work to address the ambiguities that
afflict d15Nbulk in open ocean records, with the ideal goal of quantifying diagenetic
isotope effects so that they can be removed from the bulk records. One promising
approach is the isolation and analysis of organic matter that is of known origin and
protected from diagenesis, such as microfossil-bound N. The development of
compound-specific techniques, such as for chlorophyll and amino acids, also holds
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promise for the future, as does the coupling of compound-specific techniques with
the use of microfossil-bound organic matter (Ingalls et al., 2003).
The interpretation of isotopic data will become more quantitative and mechanistic. The use of numerical models with realistic circulation, coupled with the
inclusion of other paleoceanographic proxies, will lead to much better-defined
pictures of, for example, how water column denitrification zones have interacted
with physical circulation and respiration. The inclusion of nitrogen isotope dynamics in ocean general circulation models will allow estimates of past rates of denitrification, N2 fixation and high latitude nitrate utilization to be drawn from the
expanding array of d15N records. Also, although we have limited our discussion
here to the past 2 My, the number of studies exploring the nitrogen isotopic
dynamics of ancient oceans is increasing rapidly, and is sure to reveal exciting
insights. Together, these advances will illuminate the connections of N cycle
processes to other aspects of the climate system and to each other.
This work will necessarily go hand in hand with studies to link N isotopic
patterns in the modern ocean to underlying N transformations. Some important
processes such as sedimentary denitrification can only be studied indirectly with
available or foreseeable paleoceanographic tools. These processes should be considered critical targets for improved understanding through study of modern environments. Finally, paleoceanographic techniques that, to date, have focused mostly on
open-ocean processes, can be brought to bear on the critical environment of the
coastal zone. This prospect offers a wealth of unexploited avenues of research with
implications for preindustrial human impacts on the environment, fisheries, and
coastal zone management.
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