1
Oxygen and the upside down nitrogen cycle in the Archean Ocean
Linda V. Godfrey1 and Paul G. Falkowski1,2
1
Institute of Marine and Coastal Sciences, 2Department of Earth and Planetary Sciences,
Rutgers University, New Brunswick, NJ, 08901
The isotopic composition of organic nitrogen in marine sedimentary rocks
reflects the biogeochemical history of the element at the time of deposition 1,2. We
measured the isotopic composition of nitrogen in kerogen (15Nkerogen) extracted
from minimally altered, organic-rich shales formed in the late Archean (2,700 to
2,400 Ma) 3,4 obtained from a core drilled on the Campbellrand-Malmani platform
in South Africa. Assuming that the isotopic composition of atmospheric N2 has
remained constant throughout Earth’s history 5, the samples are all slightly
enriched in 15N relative to N fixed from atmospheric N2, but there is a moderate
increase in 15N enrichment in the sediments younger than ~ 2500 Ma. The positive
values of 15N are consistent with a nitrogen cycle that requires the production of
free molecular oxygen 6,7. In the older portions of the core, there appear to be 2 to 3
cycles in 15Nkerogen, suggesting there were extended periods when oxygen
concentrations were sufficiently high in the upper ocean to allow a loss of fixed
inorganic N, either through coupled nitrification-denitrification reactions or via the
anammox reaction 8. These sequences are interspersed by one or more periods of
relatively low values of 15Nkerogen, suggesting periods of little denitrication and
extremely low oxygen concentrations. Between ~2520 and ~ 2460 Ma, 15Nkerogen
steadily approaches higher values between 2 and 4‰, suggesting the presence of a
2
stable, oxygen-dependent nitrogen cycle that allowed 15N-enriched inorganic
nitrogen to accumulate. Based on the contemporary nitrogen cycle, these results
imply that by ~ 2500 Ma, O2 in the surface ocean had crossed a concentration
threshold of about 15 M, where nitrification is competitive with denitrification 6
By far, N2 and O2 are the two most abundant gases in Earth’s atmosphere and the
fluxes of both elements are almost entirely mediated by microbes. The feedbacks
between these two elements and their fluxes constrain, in part, the oxidation state of the
planet 7. The overwhelming majority of O2 in Earth’s atmosphere is produced via the
photobiological oxidation of liquid water by oxygenic photosynthetic organisms. When
oxygenic photosynthetic organisms first arose on Earth is not known with certainty 9;
however, the pattern of isotopic fractionation of sulfur suggests that by 2300 Ma the
accumulation of O2 in the atmosphere was sufficient to allow a stratospheric ozone layer
to form, which effectively quenched UV-catalyzed, mass-independent isotopic
fractionations of that element 10,11. There is, however, compelling molecular fossil
evidence that oxygenic photosynthetic organisms (i.e., cyanobacteria) were present
several hundred million years prior to this so-called “Great Oxidation Event” (GOE) 11,12.
The question arises: “Why, given the potential for cyanobacteria to produce O2, did it
take hundreds of millions of years for the gas to accumulate in Earth’s atmosphere?” In
part, the answer appears to be related to the feedbacks between the oxygen and nitrogen
cycles.
In the contemporary nitrogen cycle, N undergoes a series of redox reactions that
produce molecules in which the oxidation state of N ranges between –III and +V.
Changes in the redox state of fixed N are sensitive to changes in the concentration of
3
oxygen. Some of these redox state changes produce the gases N2 and N2O, and are
associated with strong fractionation between the two stable isotopes of N. Molecular N2
from the atmosphere enters into biological macromolecules after reduction (“fixation”) to
NH3 through a process mediated exclusively by nitrogenase 13. Nitrogenase is, unless
protected within a cell, irreversibly inhibited by oxygen. The enzyme appears to predate
the rise of cyanobacteria and oxygen production, and was acquired by this clade via
horizontal gene transfer 14. There is virtually no isotopic fractionation associated with
nitrogen fixation. In the presence of O2, NH3 is sequentially oxidized by
chemoautotrophic bacteria to NO2- and NO3-. Under anaerobic conditions, these two
molecules can, in turn, be reduced by facultative anaerobic heterotrophs, ultimately
forming N2 15. The latter two reaction sequences are accompanied by isotopic
fractionations such that 14N is preferentially volatilized as a gaseous species, while 15N
accumulates in the dissolved pool of fixed inorganic N and can be reincorporated into
organic matter. Indeed, the coupled nitrification/denitrification can enrich the substrate
by 22 to 30‰ 2,16. Hence, changes in N isotopic composition of sedimentary organic
matter potentially reflect the oxidation state of the environment at the time of deposition
17
.
We analyzed nitrogen and organic carbon isotope ratios from drill core GKP01
that recovered more than 1300 m of Transvaal Supergroup sediment from the Griqualand
West sub-basin in South Africa (Figure 1). Here we examine the interaction of the N
cycle with the O cycle during the last 200 to 250 million years of the Archean, a critical
period in Earth’s history when several biogeochemical processes were strongly
influenced by the presence of free O2 in the upper ocean 6,18.
4
Nitrogen isotope ratios were measured on total (“bulk” N) sediments, refractory
organic molecules extracted as kerogen, and clays (“fixed” N). This apparent analytical
redundancy addresses two potential concerns: post-depositional alteration of 15N and
contamination by modern N. Regional heating in Griqualand West only reached 110 to
170º C 19; post-depositional incursion of fluids are more likely to have altered 15N 20.
The difference between bulk and kerogen or fixed 15N reflects alteration of primary
15N signals and is generally < 2‰, except in younger samples (closer to the GOE) where
it increases up to 7‰ (Figure 2). The kerogens are almost always enriched in 14N. The
difference between bulk sediment and fixed 15N is likewise small, but increases up to
4‰ in the younger samples (Supplementary Table 1). The isotopic signal preserved in
kerogens is far more likely to reflect the isotopic composition of organic matter at the
time of deposition than that of the bulk sediment. The 15N of vegetation in South Africa
varies between 2 and 7‰ 21, and may be a potential source of contamination. No modern
plant biomarkers have been found in GKP0122. Modern N that might enter the core
would be present in an inorganic form or as small molecules; the extraction of kerogen
(described in Supplementary Table 1) removes and prevents re-adsorption of surface
bound N.
We infer the effect of redox state on 15N during the late Archean using the
isotopic composition of the element in kerogen. The values of 15Nkerogen range from 0.5
to 7.3‰ across the 200 to 250 myr window recorded in the core (Figure 2). The changes
in 15N reveal 2 to 3 asymmetric peaks where 15N steps to higher values, first to 4.5 ‰
in the lower Nauga Formation (~2600 Ma) and then to ~7 ‰ in the upper portion of the
same the same Formation (~ 2520 Ma). The first stepped increase in 15N is followed by
5
a gradual decrease in 15N towards 0.5 ‰, and is preceded by a decline in 15N from an
earlier high value of 15N that predates the lowermost section of the core. Following the
rise in 15N within the upper Nauga Formation, 15N appears to trend towards an
asymptote of ~ 2 to 4 ‰ within the Klein Naute Formation which spans 2520 to 2460
Ma.
GKP01’s location on the slope of the Campbellrand-Malmani carbonate platform
mostly exceeded storm weather wave base 23, while the overwhelming proportion of
organic matter deposited in the sediments was formed in the overlying euphotic zone
ocean. The isotopic values preserved in the sediments integrate the nitrogen sources that
fueled the exported flux from the upper ocean. Phototrophic organisms in the surface
ocean could have utilized two basic forms of fixed inorganic N: (a) “regenerated” N,
derived from biologically recycling in the surface ocean, (b) “new” N in the form of
NH4+ , that was either upwelled from the ocean interior or fixed from atmospheric N2
(Figure 3). On geological time scales, the sinking flux is proportional to the flux of the
latter. Although ultimately all forms of N assimilated by organisms are incorporated into
organic matter as NH3, the presence of oxygen in upper ocean zone would have made two
sources isotopically distinct. Through coupled nitrification-denitrification of fixed N in
the surface ocean and preferential loss of 14N, the 15N of regenerated N would have been
higher than “new” N, which initially would have been isotopically very similar to the
atmospheric 15NN2 of 0. Following this logic, changes in 15N reflect the mixing ratio of
two end member sources: the proportion of fixed-N assimilated that had been partially
nitrified in the upper ocean versus that which upwelled in the form of NH4+. Hence, the
“new” nitrogen in the Archean ocean, supplied from the interior was a reduced species,
6
while the “regenerated” nitrogen in the surface waters was oxidized. In effect, the
Archean ocean was “upside-down” relative to modern basins 24.
To first order, the 15N of organic matter reflects the relative proportion of
production supported by the two N sources. The changes in 15N measured in GKP01
show that neither the isotope composition nor inventory of fixed N in the ocean remained
constant during the late Archean. During the deposition of GKP01 from ~2700 Ma to
~2460 Ma, there are two, and possibly three, cyclical changes in 15N suggesting that
there were at least two to three times when oxygen production and coupled nitrificationdenitrification led to a loss of ocean fixed- N. Since the Archean oceanic N isotope
budget was determined by denitrification sinks and N-fixation sources (biological and
abiotic), the return of 15N values close to 0 indicates that rates of N-fixation were able
resupply lost N, albeit at a rate much slower than the rate at which the ocean was
denitrified. Further, assuming constant rates of N-fixation, each of these cycles of N loss
and gain were directly linked to increases and decreases in oxygen production.
Progressively higher rates of O production and N loss resulted in slightly higher 15N for
each cycle (Figure 2). That very high values of 15N are not recorded suggests that at no
time did the ocean become completely denitrified.
Between 2520 and 2460 million years ago, the values of 15N began to trend
upwards towards values between 2 and 4‰ (Figure 2). In order to maintain these
positive 15N values, simple partial nitrification-denitrification of N added to the ocean
would lead to an indefinite increase in the fixed N inventory unless there was a way to
remove N with minimal isotope effect. In the modern ocean this can be achieved by
denitrification of nitrate in deep ocean sediments 25; however, in an ocean where NH4+
7
was the dominant form of N in the interior, such a scenario is highly improbable. The
most likely scenario is that oxygen production had allowed the concentration of oxygen
in the surface ocean to reach a threshold of ~15 M, where the rate of nitrification can
exceed denitrification 6. Denitrification of this reservoir of nitrate led to its strong
enrichment of 15N that could then have been available for assimilation, recycling through
organic matter and export to the deep ocean.
When oxygenic photosynthetic organisms first arose on Earth remains unclear 9.
Initial estimates of 2800 to 2700 Ma , based on hopanoid biomarkers 26, have been
questioned as these molecules can also be produced by anoxygenic phototrophs 27.
However, it is possible that oxygen production evolved well before the loss of sulfur
mass independent isotopic fractionation signals or appearance of redbeds 10,11,28 imply the
existence of free oxygen in the geological record. Indeed, the Nitrogen isotope
compositions of organic matter produced in the surface of the Archean ocean indicate
that oxygen production led to accumulation of small concentrations of free molecular
oxygen several times towards the end of the Archean. Oxygenic photosynthetic
organisms may have taken advantage of the huge epeiric seas that covered the Kaapvaal
and Pilbara cratons between 2600 and 2400 Ma 29 which, in todays ocean, are regions of
high biological activity driven by coastal upwelling and efficient nutrient recycling.
Through coupled nitrification-denitrification, oxygen production led to reductions in the
N inventory; these periods were characterized by increasing 15N. Positive 15N values
that were maintained toward the end of the Archean occurred as pO2 crossed a threshold
allowing nitrification to outpace denitrification, and 15N-enriched nitrate to accumulate.
8
Figure 1 Map showing the distribution of the Campbellrand-Malmani platform on the
Kaapvaal craton in South Africa, and the location of the Agouron Drilling Project core
GKP01 in the Griqualand West sub-basin.
Figure 2 Changes in the abundance of organic nitrogen and carbon and the isotopic
composition of nitrogen in bulk sediments and the kerogen fraction over the last 200
million years of the Archean recorded in core GKP01. The isotopic composition is
correlated with water depth up to ~ 580 m in the core. Above 580 m 15Nkerogen continues
to increase due to an increase in accumulation of oxidants, including nitrate due to
increased organic carbon burial leading to decoupling between photosynthesis and
respiration. The difference between kerogen and bulk sediment 15N is most likely due
to isotope fractionation of loosely bound N that entered the core after deposition, possibly
derived from fluids associated with a hydrothermal event about 2,200 Ma 30. There is no
correlation between 15Nkerogen and 13Corg (refer to on-line supplemental information)
indicating that thermal maturation has is not responsible for the pattern of variation in
15Nkerogen. Dates for the units are: Ventersdorp Supergroup 2710 – 2700 Ma 31; lower
Nauga Fm. 2588 ± 6 Ma 32; upper Nauga above Kamden Mb. 2552 ± 11 Ma and 2549 ± 7
Ma, 4,33; upper Nauga below Klein Naute Fm. 2521 ± 3 3; Kuruman Fm. 2460 ± 5 Ma 34.
Figure 3 Hypothesized nitrogen cycle for the late Archean. Upwelling of ocean interior
NH4+, N-fixation and efficient N recycling support surface productivity. Oxygenenic
9
photosynthesis leads to coupled nitrification and denitrification which causes an increase
in 15N and decrease in N-inventory. At the end of the Archean, rates of oxygen
production increase allowing nitrification rates to exceed those of denitrification.
Partially denitrified nitrate with high 15N can accumulate and be available for
assimilation. Eventually the isotope signature of this process is carried by sinking
organic matter to the ocean interior.
Acknowledgements: This research was supported by a grant from the Agouron
Foundation. Additional support was provided by NASA. We thank Woodward Fischer,
Andy Knoll, Ed Stolper, Nic Beukes and Joe Kirschvink for comments.
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12
Godfrey and Falkowski, Figure 2
13
Supplementary Table. Isotope data for Agouron Drilling Project core GKP01
Sample depth (m)
15Nbulk
197.93-198.05
244.69
267.18
279.39
292.2
310.57
312.83
322.12 - 322.22
360.7
413.63
482.69 - 482.81
491.15
494.2
499.32
515.95
574.76
606.5-606.59
614.55
626
637
637.70 - 637.81
640.3
682.52-682.6
689.1
721.63
733.4
738.95
743.1
798.46 - 798.51
798.5
871.83
890.60 - 890.68
891.25
892.4
965.1
970.75
974.05
980.04-980.13
1008.00 - 1008.05
1012.1
1019.9
1029.01 - 1029.10
1035.7
10.09
4.41
10.12
7.61
4.59
5.25
4.56
4.29
4.54
6.72
3.57
3.47
4.16
7.10
4.42
8.50
5.65
4.06
7.97
2.41
3.37
2.45
5.45
2.13
1.75
2.37
3.87
5.76
3.70
3.31
4.71
3.98
8.53
5.93
2.87
4.13
5.79
3.31
4.66
1.95
3.27
2.35
15N
kerogen a
15N fixed b
3.11
4.01
5.00
2.83
4.50
2.17
3.20
2.69
6.03
7.27
2.84
5.61
3.93
3.27
4.29
4.07
6.47
6.91
3.91
0.56
3.93
0.54
3.72
3.1
2.19
4.46
-1.25
3.00
2.62
3.25
2.69
3.20
2.95
4.04
4.19
2.53
0.92
2.30
0.50
2.58
13C
organic c
-33.52
-32.76
-33.79
-34.62
-38.73
-36.05
-32.58
-30.19
-33.32
-32.24
-31.79
-32.79
-28.68
-31.48
-33.35
-31.94
-31.64
-38.24
-38.23
-38.56
-28.14
-33.78
-32.51
-30.52
-32.2
-30.52
-29.89
-37.53
-35.18
-28.68
-20.38
-25.04
-35.19
-35.69
-33.14
-36.45
-32.45
-37.23
-36.92
-36.3
14
1036.8-1036.87
1047.7
1077
1079.9
1087.8
1150.9
1155.15
1163
1200.95
1272.1
1294.8
3.11
3.41
3.75
3.25
2.60
3.25
3.88
4.52
4.35
5.48
2.33
2.51
3.49
-38.95
-35.07
-35.27
-34.84
-34.93
-38.05
-35.87
-31.83
-42.42
-37.8
-31.39
Sample fragments, which did not include any cut edges, were powdered and used for all
analyses. Nitrogen and C isotope compositions were measured using a Eurovector
Elemental Analyser coupled to a GVI Isoprime mass spectrometer at Rutgers University,
except for data in grey which were prepared by off-line combustion and analyzed using a
dual inlet ThermoFinnigan 252 mass spectrometer at Lehigh University. Samples
analyzed by CF-IRMS had V2O5 added to aid combustion. Multiple blank filters and
certified standards, IAEA N1, IAEA N3, and NBS 22, were included in each run and to
which the data were corrected. Standard deviations for 13C and 15N on triplicate
samples were 0.15‰ and <0.3‰ for bulk sediment; 0.2‰ for fixed N and <0.4‰ for
kerogen N.
a
Kerogen was extracted from powdered samples by digesting 200 mg decalcified
sediment in 20 mL 3N HCl – 15N HF at 50ºC. The samples were centrifuged, the acid
poured off and then rinsed with 0.5 N KCl solution three times to remove remaining acid.
The samples were then filtered on to precombusted GF/F filters and analyzed following
the described method above.
b
Nitrogen fixed in clay minerals was analyzed by removing organic bound and surface
adsorbed N from decalcified samples and then following the method of Silva and
Bremner (1966)1.
c
Samples were decarbonated in silver capsules using 10% HCl.
1. Silva, J. A., Bremner, J. M. Determination and isotope-ratio analysis of different forms
of nitrogen in soil: 5. Fixed ammonium. Proceedings of the Soil Science Society of
America 30 (1966) 587-594.