Introduction
Our ability to mitigate soil N2O emissions is limited due to poor
understanding of the complex interplay between N2O production pathways
in soil environments. In order to develop effective fertilizing strategies
and reduce the loss of nitrogen through microbial consumption as well as
related adverse environmental impacts (IPCC, 2013; Ravishankara et al.,
2009), it is very important to fill the existing knowledge gaps. Isotopocule
analyses of N2O, including δ18O, average δ15N
(δ15Nav) and 15N site preference within the linear
N2O molecule (δ15Nsp) have been used for several
years to help differentiate between N2O production pathways (Opdyke et
al., 2009; Perez et al., 2006; Sutka et al., 2006; Toyoda et al., 2005; Well
et al., 2008), the various microbes involved (Rohe et al., 2014a; Sutka et
al., 2003, 2008) and to estimate the fraction of N2O reduced to N2
(Ostrom et al., 2007; Park et al., 2011; Toyoda et al., 2011; Well and
Flessa, 2009). However, the usefulness of these analyses would be enhanced
further if the isotope fractionation mechanisms were better understood. In
particular, we need to recognize the isotope effects associated with nitrate
and N2O reduction to quantify the entire gaseous nitrogen losses as
N2O and N2 based on the N2O isotopic signatures
(Lewicka-Szczebak et al., 2014, 2015). This would be most effective if either
of the isotopic signatures (δ18O, δ15Nav or
δ15Nsp) were stable or predictable for N2O produced
by each of the relevant N2O forming processes (e.g. heterotrophic
bacterial denitrification, fungal denitrification, nitrifier denitrification
and nitrification). We hypothesize that this could be the case for δ18O, and this study aims to increase the understanding of the factors
controlling δ18O during N2O production in soils.
δ18O(N2O) has been rarely applied in the interpretation of
N2O isotopic signatures because of the rather complex oxygen isotope
fractionations during N2O production by denitrification (Kool et al.,
2007). Denitrification is a stepwise process of nitrate reduction mediated by
three enzymes: nitrate reductase (NAR), nitrite reductase (NIR) and nitric
oxide reductase (NOR) (Fig. 1). δ18O(N2O) is controlled by the
origin of the oxygen atom in the N2O molecule (nitrate, nitrite, soil
water or molecular O2) and by the isotope fractionation during nitrate
reduction or during oxygen isotope exchange with soil water.
Oxygen isotope fractionation during denitrification as a result of
branching effects
(εNAR,εNIR,εNOR) and
exchange effects (εw) associated with the following
enzymatic reaction steps: NAR, NIR and NOR.
During each reduction step, one oxygen atom is detached and removed as water
(H2O), which is associated with branching isotope effects (Casciotti et
al., 2007; Snider et al., 2013). Conceptually, these can be regarded as a
combination of two isotope fractionations with opposite effects on the
δ18O signature of the reduction product: (i) intermolecular
fractionation due to preferential reduction of 18O-depleted molecules,
which results in 18O-enriched residual substrate and 18O-depleted
product, and (ii) intramolecular fractionation due to preferential 16O
abstraction, which results in 18O-enriched nitrogen-bearing reduction
products and 18O-depleted H2O as side product. Since intermolecular
fractionation causes 18O depletion of the reduction product and
intramolecular fractionation causes 18O enrichment, the net branching
effect (εn), as the sum of both, can theoretically vary
between negative and positive values. However, pure cultures studies show
that εn is mostly positive, i.e. between 25 and
30 ‰ for bacterial denitrification (Casciotti et al., 2007) and
between 10 and 30 ‰ for fungal denitrification (Rohe et al., 2014a).
Importantly, the intra- and intermolecular isotope effects can only manifest
together during incomplete substrate consumption (Rohe et al., 2014a). In the
case of complete substrate conversion, the net branching effect reflects the
intramolecular effect only (Casciotti et al., 2007).
Moreover, denitrification intermediates may partially or fully exchange
oxygen isotopes with ambient water (Kool et al., 2009). The isotopic
signature of the incorporated O-atom depends on the isotopic signature of
ambient water and the isotope fractionation associated with this exchange.
Under typical soil conditions, i.e. pH close to neutral and moderate
temperatures, abiotic isotope exchange between nitrate and water is
negligibly slow. In extremely acid conditions (pH < 0), the
equilibrium effect is ε(NO3- / H2O) = 23 ‰ (Böhlke et al., 2003).
Casciotti et al. (2007) showed that for nitrite the abiotic exchange can
occur at neutral pH, but for achieving an isotopic equilibrium over 8 months
are needed. The observed isotope equilibrium effect between nitrite and water
is ε(NO2- / H2O) = 14 ‰ at
21 ∘C. Nothing is known yet about the possible abiotic exchange
between NO and ambient water. The isotope exchange between denitrification
intermediates and ambient water is most probably accelerated by enzymatic
catalysis, since numerous 18O tracer studies documented nearly complete
O isotope exchange (Kool et al., 2009; Rohe et al., 2014b; Snider et al.,
2013) within short incubation times like a few hours. Hence, it can be
assumed that at least one enzymatic step must be responsible for exchange of
O isotopes with soil water (Rohe et al., 2014a; Snider et al., 2013). In pure
culture studies the extent of oxygen isotope exchange ranged from 4 to
100 % for bacterial denitrification (Kool et al., 2007) and from 11 to
100 % for fungal denitrification (Rohe et al., 2014b). In contrast,
unsaturated soil incubation experiments, with a natural whole microbial
community, showed consistently high magnitudes of oxygen isotope exchange of
between 85 and 99 % (Kool et al., 2009; Lewicka-Szczebak et al., 2014;
Snider et al., 2013). If the high extent of isotope exchange was
characteristic of soil denitrification processes, we would expect quite
stable δ18O values of the produced N2O during denitrification.
It is difficult to quantitatively link isotope exchange and apparent isotope
effects, because using the 18O tracer technique to quantify isotope
exchange prevents simultaneous study of isotope oxygen fractionation.
However, two studies that conducted parallel 18O traced and natural
abundance experiments allowed formulating general oxygen isotope
fractionation models (Rohe et al., 2014a; Snider et al., 2013). These models
showed that the magnitude of overall isotope fractionation depends not only
on the extent of oxygen isotope exchange but also on the enzymatic reduction
step associated with this exchange (Fig. 1). It was found that the oxygen
isotope exchange is predominantly associated with NIR for fungal
denitrification (Rohe et al., 2014a). Fungi and bacteria are characterized by
different NOR mechanisms (Schmidt et al., 2004; Stein and Yung, 2003),
resulting in distinct δ15Nsp values for bacterial and
fungal denitrification. It is possible that these different NOR mechanisms
also influence δ18O.
In the present study, we used 17O as tracer to determine the extent of O
isotope exchange, in order to investigate both isotope exchange and apparent isotope
effects. We applied a nitrate fertilizer of natural atmospheric deposition
origin with high 17O excess, as a result of non-random oxygen isotope
distribution. Then we measured 17O excess of the produced N2O and,
based on the observed loss of 17O excess, calculated the extent of
isotope exchange with water. Simultaneously, we could measure the
18O / 16O fractionation in the same incubation vessels, since
the 17O tracing method has no impact on δ18O. This is the
first time that such an approach has been used. To validate this method, we
applied an alternative approach – namely, soil water with distinct
δ18O values within the range of natural abundance isotopic
signatures was applied to quantify isotope exchange (Snider et al., 2009).
The latter method has also been applied in a recent soil incubation study
(Lewicka-Szczebak et al., 2014) and indicated almost complete oxygen isotope
exchange with soil water associated with a stable isotope ratio difference
between soil water and produced N2O of
δ18O(N2O / H2O) = (19.0 ± 0.7) ‰.
However, the results of other experiments presented in the same study
(Lewicka-Szczebak et al., 2014) indicated much higher
δ18O(N2O / H2O) values of up to 42 ‰. The
higher values may be due to a lower extent of oxygen isotope exchange, but no
data were available regarding the extent of exchange for those samples. In
the present study, we investigated possible controlling factors for oxygen
isotope exchange by applying various experimental treatments differing in
soil moisture and temperature.
The combination of various experimental approaches allowed us to further
improve the δ18O fractionation model proposed by Snider et
al. (2013) and Rohe et al. (2014a), to decipher the mechanism of oxygen
isotope fractionation during N2O production by denitrification and to
determine the associated isotope effects. We investigated the variability of
isotope exchange with soil water and of the δ18O values of produced
N2O under varying conditions as well as the relation between these
quantities. Ultimately, our aim was to check to what level of accuracy
δ18O can be predicted based on the known controlling factors.
Additionally, the 17O analyses of N2O produced by denitrification
gave us the opportunity to test the hypothesis of soil denitrification
contributing to the non-random distribution of oxygen isotopes (17O
excess, or Δ17O) in atmospheric N2O (Kaiser et al.,
2004; Michalski et al., 2003).
Methods
Experimental set-ups
Experiment 1 (Exp 1) – static anoxic incubation
The static incubations were performed under an anoxic atmosphere (N2) in
closed, gas-tight vessels where denitrification products accumulated in the
headspace. Two arable soil types were used: a Luvisol with loamy
sand texture and Haplic Luvisol with silt loam texture with
pH (in 0.01 M CaCl2) of 5.7 and 7.4, respectively. More details on soil
properties can be found in Lewicka-Szczebak et al. (2014). For the first part
of these incubations (Exp 1.1) two different temperature treatments were
applied (8 and 22 ∘C) and only one moisture treatment of 80 %
WFPS (water-filled pore space). The results of δ18O(N2O)
analyses for these samples have already been published (Lewicka-Szczebak et
al., 2014). Here we expand these data with Δ17O(N2O) analyses.
The second part of the static incubations (Exp 1.2) was performed for the
same two soils with three different moisture treatments of 50, 65 and
80 % WFPS at one temperature (22 ∘C). Details on the treatments
are presented as supplementary information in Supplement Table S1.
This experimental approach is described in detail in Lewicka-Szczebak et
al. (2014). In short, the soil was air dried and sieved at 2 mm mesh size.
Afterwards, the soil was rewetted to obtain the target WFPS and fertilized
with 50 (Exp 1.1) or 10 (Exp 1.2) mg N equivalents (as NaNO3) per kg
soil. Various nitrate and water treatments were applied (Table S1). The soils
were rewetted using two waters with distinct isotopic signatures –
heavy water (δ18O =-1.5 ‰) and light water (δ18O = -14.8 ‰) – and fertilized with two
different nitrate fertilizers – natural Chile saltpeter (NaNO3,
Chili Borium Plus, Prills-Natural origin, supplied by Yara, Dülmen,
Germany, δ18O =56 ‰) and synthetic NaNO3
(Sigma Aldrich, Taufkirchen, Germany, δ18O =27 ‰). The
soils were thoroughly mixed to obtain a homogeneous distribution of water and
fertilizer and an equivalent of 100 g of dry soil was repacked into each
incubation jar at bulk densities of 1.3 g cm-3 for the silt loam soil
and 1.6 g cm-3 for the loamy sand soil. The 0.8 dm3 jars
(J. WECK GmbH u. Co. KG, Wehr, Germany) were used with airtight rubber seals
and with two three-way valves installed in their glass cover to enable
sampling and flushing. The jars were flushed with N2 at approximately
500 cm3 min-1 (STP: 273.15 K, 100 kPa) for 10 min to create
anoxic conditions. Immediately after flushing, acetylene (C2H2) was
added to inhibit N2O reduction in selected jars (C2H2
inhibited treatment), by replacing 80 cm3 of N2 with
C2H2, which resulted in 10 kPa C2H2 in the headspace.
Each treatment (Table 1) had three replicates. The soils were incubated for
approximately 25 h and three to four samples were collected at 4–12 h intervals by transferring 30 cm3 of headspace gases into two
pre-evacuated 12 cm3 Exetainer vials (Labco Limited, Ceredigion, UK).
The excess 3 cm3 of headspace gas in each vial ensured that no ambient
air entered the vials. The removed sample volume was immediately replaced by
pure N2 gas.
Exp 1 results: soil moisture (expressed as water-filled pore space:
WFPS), N2O + N2 production rate (expressed as mass of N as sum
of N2O and N2 per mass of dry soil per time), 17O excess in
soil nitrate (Δ17O(NO3)) and in N2O (Δ17O(N2O)) with calculated exchange with soil water (x), and oxygen
isotopic signature (δ18O) of soil nitrate (NO3-), soil
water (H2O) and N2O with calculated isotope ratio difference
between soil water and N2O (δ018O
(N2O / H2O)). For samples with non-inhibited N2O reduction
the N2O mole fraction (f(N2O)) was taken into account to
calculate the δ18O unaffected by N2O reduction (δ018O(N2O)) and the respective δ018O(N2O / H2O). Only Chile saltpeter treatments are
presented, for which the individual determination of x was possible. Part
of the data from Exp 1.1 (δ18O(NO3-), δ18O(H2O), δ18O(N2O)) has already been published in
Lewicka-Szczebak et al. (2014).
treatment
N2O + N2
Δ17O(NO3-)
Δ17O(N2O)
x (%)
δ18O(NO3)
δ18O(H2O)
δ18O(N2O)
f(N2O)a
δ018O (N2O)b
δ018O (N2O / H2O)
production rate
(µg kg-1 h-1)
(‰)
(‰)
(‰)
(‰)
(‰)
(‰)
(‰)
WFPS
inhibition
(%)
Exp 1.1 a, loamy sand, 8 ∘C
80
114
11.9 ± 0.6
0.4 ± 0.5
96.2 ± 4.7
38.8 ± 0.5
-9.2 ± 0.5
13.4 ± 0.2
0.84 ± 0.04
10.4
19.7 ± 0.5
80
C2H2
107
11.9 ± 0.6
0.8 ± 0.4
93.1 ± 3.1
38.8 ± 0.5
-9.2 ± 0.5
10.4 ± 0.1
1
10.4
19.8 ± 0.5
80
125
11.9 ± 0.6
0.8 ± 0.2
92.7 ± 1.1
37.5 ± 0.5
-13.5 ± 0.5
8.4 ± 0.3
0.84 ± 0.04
5.4
19.1 ± 0.6
80
C2H2
126
11.9 ± 0.6
0.3 ± 0.7
96.2 ± 3.4
37.5 ± 0.5
-13.5 ± 0.5
5.7 ± 0.0
1
5.7
19.4 ± 0.5
Exp 1.1b, loamy sand, 22 ∘C
80
427
10.4 ± 0.8
0.4 ± 0.2
95.7 ± 1.8
42.6 ± 0.5
-9.2 ± 0.5
12.5 ± 0.2
0.85 ± 0.06
9.6
19.0 ± 0.5
80
C2H2
362
10.4 ± 0.8
0.4 ± 0.0
96.4 ± 0.2
42.6 ± 0.5
-9.2 ± 0.5
9.5 ± 0.0
1
9.5
18.9 ± 0.5
80
429
10.4 ± 0.8
0.2 ± 0.1
98.2 ± 1.5
42.1 ± 0.5
-13.5 ± 0.5
7.5 ± 0.1
0.85 ± 0.06
4.7
18.4 ± 0.5
80
C2H2
370
10.4 ± 0.8
0.5 ± 0.1
94.8 ± 0.5
42.1 ± 0.5
-13.5 ± 0.5
4.5 ± 0.1
1
4.5
18.3 ± 0.5
Exp 1.1 c, silt loam, 22 ∘C
80
266
9.2 ± 1.3
0.0 ± 0.2
99.5 ± 0.9
31.8 ± 0.5
-2.6 ± 0.5
26.4 ± 0.1
0.57 ± 0.03
16.4
19.1 ± 0.5
80
C2H2
257
9.2 ± 1.3
0.4 ± 0.1
95.3 ± 1.4
31.8 ± 0.5
-2.6 ± 0.5
15.9 ± 0.1
1
15.9
18.5 ± 0.5
80
271
9.2 ± 1.3
0.1 ± 0.2
98.6 ± 1.3
31.8 ± 0.5
-8.7 ± 0.5
20.7 ± 0.2
0.57 ± 0.03
10.8
19.7 ± 0.5
80
C2H2
251
9.2 ± 1.3
0.4 ± 0.1
95.0 ± 1.5
31.8 ± 0.5
-8.7 ± 0.5
9.8 ± 0.1
1
9.8
18.7 ± 0.5
Exp 1.2 a, loamy sand, 22 ∘C
80
C2H2
126
3.4 ± 0.5
n.d.
n.d.
6.5 ± 0.5
-10.4 ± 0.5
6.3 ± 0.1
1
6.3
16.9 ± 0.5
65
C2H2
112
3.4 ± 0.5
0.2 ± 0.3
92.6 ± 8.5
6.5 ± 0.5
-10.1 ± 0.5
6.9 ± 0.2
1
6.9
17.2 ± 0.5
50
C2H2
50
3.4 ± 0.5
0.0 ± 0.3
95.8 ± 3.9
6.5 ± 0.5
-8.9 ± 0.5
7.6 ± 0.3
1
7.6
16.6 ± 0.6
80
C2H2
161
3.4 ± 0.5
n.d.
n.d.
6.5 ± 0.5
-5.0 ± 0.5
10.5 ± 0.0
1
10.5
15.6 ± 0.5
65
C2H2
102
3.4 ± 0.5
0.2 ± 0.2
92.7 ± 5.2
6.5 ± 0.5
-5.7 ± 0.5
11.6 ± 0.1
1
11.6
17.5 ± 0.5
50
C2H2
74
3.4 ± 0.5
0.2 ± 0.2
94.5 ± 5.1
6.5 ± 0.5
-6.6 ± 0.5
10.7 ± 0.1
1
10.7
17.4 ± 0.5
Exp 1.2 b, silt loam, 22 ∘C
80
C2H2
137
2.6 ± 0.4
0.2 ± 0.2
90.6 ± 7.3
3.2 ± 0.5
-8.1 ± 0.5
8.3 ± 0.1
1
8.3
16.5 ± 0.5
65
C2H2
130
2.6 ± 0.4
0.2 ± 0.1
92.2 ± 3.7
3.2 ± 0.5
-7.1 ± 0.5
9.8 ± 0.1
1
9.8
17.1 ± 0.5
50
C2H2
121
2.6 ± 0.4
0.1 ± 0.1
96.5 ± 4.3
3.2 ± 0.5
-5.9 ± 0.5
12.5 ± 0.2
1
12.5
18.6 ± 0.5
80
C2H2
111
2.6 ± 0.4
-0.1 ± 0.1
99.1 ± 1.6
3.2 ± 0.5
-1.6 ± 0.5
15.1 ± 0.2
1
15.1
16.7 ± 0.6
65
C2H2
132
2.6 ± 0.4
0.0 ± 0.1
98.4 ± 1.6
3.2 ± 0.5
-1.8 ± 0.5
15.2 ± 0.2
1
15.2
17.0 ± 0.5
50
C2H2
106
2.6 ± 0.4
-0.2 ± 0.0
100.0 ± 1.8
3.2 ± 0.5
-2.0 ± 0.5
15.7 ± 0.3
1
15.7
17.7 ± 0.6
a c(N2O) / [c(N2)+c (N2O)]: based on parallel
15N treatment (last sampling results).
Where b N2O reduction not inhibited, the values are corrected taking
into account product ratio and isotope fractionation, according to Rayleigh
fractionation 18ε(N2 / N2O) values taken from
Lewicka-Szczebak et al. (2014): -17.4 ‰ (see Sect. 2.5 for
details).
Additional treatments with addition of 15N-labelled NaNO3 (98 %
15N isotopic purity) were used to control the efficiency of acetylene
inhibition and to determine the N2O mole fraction
f(N2O) = c(N2O) / [c(N2)+c(N2O)] (c: volumetric
concentration) in non-inhibited treatments. This method allows determination
of the N2 concentration originating from the 15N labelled pool and
hence the N2O mole fraction (Lewicka-Szczebak et al., 2013).
Experiment 2 (Exp 2) – flow-through incubation under He
atmosphere
The flow-through incubations were performed using a special gas-tight
incubation system allowing for incubation under N2-free atmosphere to
enable direct quantification of soil N2 fluxes (Butterbach-Bahl et al.,
2002; Scholefield et al., 1997). This system has been described in detail by
Eickenscheidt et al. (2014). Four different soils were incubated: two arable
soils, the same as in Exp 1 (loamy sand and silt loam), and two grassland soils:
an organic soil classified as Histic Gleysol and a sandy soil
classified as Plaggic Anthrosol, with pH (in 0.01 M CaCl2) of
5.9 and 5.3, respectively. All soils were incubated at the target moisture
level of 80 % WFPS and the two most active soils (organic and silt loam
soil) were additionally incubated at the lower moisture level of 70 %
WFPS (target values, for actual values see Table 2).
Exp 2 results: soil moisture (expressed as water-filled pore
space: WFPS), N2O+N2 production rate (expressed as mass of N as
sum of N2O and N2 per mass of dry soil per time), 17O excess
in soil nitrate (Δ17O(NO3)) and in N2O (Δ17O(N2O)) with calculated exchange with soil water (x) and oxygen
isotopic signature (δ18O) of soil nitrate (NO3), soil water
(H2O) and N2O. All δ18O(N2O) values were corrected
taking into account N2O mole fraction (f(N2O)) to calculate the
values unaffected by N2O reduction (δ018O(N2O)) and
the respective δ018O(N2O / H2O).
WFPS (%)
N2O+N2
Δ17O(NO3-)
Δ17O(N2O)
x (%)
δ18O(NO3-)
δ18O(H2O)
δ18O(N2O)
f(N2O)a
δ018O (N2O)b
δ018O (N2O / H2O)
production rate
(µg kg-1 h-1)
(‰)
(‰)
(‰)
(‰)
(‰)
(‰)
(‰)
Exp 2.1, sand
73.6 ± 0.7
91
10.8 ± 0.3
2.7 ± 0.4
73.9 ± 4.2
34.3 ± 1.7
-8.6 ± 0.5
12.1 ± 0.2
0.95 ± 0.01
11.5 ± 0.2
20.2 ± 0.5
2.6 ± 1.1
74.4 ± 11.0
11.0 ± 0.4
0.92 ± 0.01
10.0 ± 0.5
18.8 ± 0.7
Exp 2.2 loamy sand
70.4 ± 0.9
49
11.9 ± 0.3
3.7 ± 0.4
66.9 ± 3.1
43.0 ± 2.4
-7.4 ± 0.5
18.4 ± 2.7
0.80 ± 0.05
15.7 ± 2.1
23.3 ± 2.2
3.3 ± 0.2
71.2 ± 1.6
15.7 ± 0.9
0.83 ± 0.02
13.5 ± 0.7
21.0 ± 0.8
Exp 2.3 silt loam
78.4 ± 1.9
80
11.3 ± 0.2
5.2 ± 0.2
52.0 ± 2.2
43.1 ± 2.3
-5.3 ± 0.5
43.8 ± 2.2
0.32 ± 0.03
29.4 ± 2.6
34.9 ± 2.6
5.3 ± 0.1
50.4 ± 1.4
46.1 ± 3.9
0.29 ± 0.10
30.4 ± 0.2
35.9 ± 0.5
Exp 2.4 silt loam
73.6 ± 1.8
52
12.1 ± 0.3
3.5 ± 0.5
69.9 ± 4.0
52.0 ± 3.3
-5.0 ± 0.5
30.1 ± 0.4
0.68 ± 0.02
25.4 ± 0.7
30.5 ± 0.9
5.0 ± 0.5
56.3 ± 4.1
37.7 ± 4.1
0.63 ± 0.07
31.9 ± 4.3
37.1 ± 4.3
Exp 2.5 organic
86.5 ± 1.8
743
7.8 ± 0.2
2.3 ± 1.1
68.1 ± 13.8
30.4 ± 0.6
-6.4 ± 0.5
26.4 ± 5.3
0.60 ± 0.02
20.0 ± 5.1
26.6 ± 5.1
2.3 ± 0.8
68.2 ± 9.5
37.7 ± 2.9
0.51 ± 0.02
29.3 ± 3.3
36.0 ± 3.3
Exp 2.6 organic
78.7 ± 0.4
1198
12.5 ± 0.7
1.1 ± 0.2
90.2 ± 1.8
43.6 ± 5.6
-6.7 ± 0.5
18.5 ± 0.0
0.82 ± 0.02
16.1 ± 0.2
22.9 ± 0.6
2.3 ± 0.3
78.8 ± 3.0
25.6 ± 0.8
0.74 ± 0.05
21.9 ± 1.6
28.7 ± 1.7
ac(N2O) / [c(N2)+c(N2O)]: based on direct GC
measurements in N2-free atmosphere.
b initial δ18O values of unreduced N2O calculated
according to Rayleigh fractionation, 18ε(N2 / N2O)
values taken from Lewicka-Szczebak et al. (2015): -12 ‰ (see
Sect. 2.5).
The soils were air dried and sieved at 4 mm mesh size. Afterwards, the soil
was rewetted to obtain 70 % WFPS and fertilized with 50 mg N equivalents
(as NaNO3) per kg soil with natural fertilizer Chile saltpeter.
The soils were thoroughly mixed to obtain a homogeneous distribution of water
and fertilizer and 250 cm3 of wet soil was repacked into each
incubation vessel at bulk densities of 1.4 g cm-3 for the silt loam
soil, 1.6 g cm-3 for the loamy sand soil, 1.5 g cm-3 for the
sandy soil, and 0.4 g cm-3 for the organic soil. Afterwards, the water
deficit to the target WFPS was added on the top of the soil for 80 % WFPS
treatments. Each treatment had three replicates. The incubation vessels were
cooled to 2 ∘C and repeatedly evacuated (to 4.7 kPa) and flushed
with He to reduce the N2 background and afterwards flushed with a
continuous flow of 20 % O2 in helium (He / O2) mixture at
15 cm3 min-1 (STP) for at least 60 h. When a stable and low-N2 background (below 10 µmol mol-1) was reached,
temperature was increased to 22 ∘C. During the incubation the
headspace was constantly flushed with He / O2 mixture (first 3 days;
Part 1) and then with He (last 2 days; Part 2) at a flow rate of
approximately 15 cm3 min-1 (STP). The fluxes of N2O and
N2 were analysed immediately (see Sect. 2.2) and f(N2O) was
determined. Samples for N2O isotopocule analyses were collected by
connecting the sampling vials in line with the exhaust gas of each incubation
vessels and exchanging them at least twice a day. The results presented in
this study originate from the anoxic Part 2 of the incubation, since the
N2O fluxes during the Part 1 were too low for Δ17O analyses.
The results for two samples taken approximately 8 and 24 h after switch to
anoxic conditions are shown.
Gas chromatographic analyses
In Exp 1 the samples for gas concentration analyses were collected in
Exetainer vials (Labco Limited, Ceredigion, UK) and were analysed using an
Agilent 7890A gas chromatograph (GC) (Agilent Technologies, Santa Clara, CA,
USA) equipped with an electron capture detector (ECD). Measurement
repeatability as given by the relative standard deviation (1σ) of
four standard gas mixtures was typically 1.5 %.
In Exp 2, online trace gas concentration analysis of N2 was performed
with a micro-GC (Agilent Technologies, 3000 Micro GC), equipped with a
thermal conductivity detector (TCD) and N2O was measured with a GC
(Shimadzu, Duisburg, Germany, GC–14B) equipped with ECD detector. The
measurement repeatability (1σ) was better than 0.02 for N2O and
0.2 µmol mol-1 for N2.
Isotopic analyses
Isotopocules of N2O
Gas samples were analysed using a Delta V isotope ratio mass spectrometer
(Thermo Scientific, Bremen, Germany) coupled to automatic preparation system:
Precon + Trace GC Isolink (Thermo Scientific, Bremen, Germany) where
N2O was preconcentrated, separated and purified. In the mass
spectrometer, N2O isotopocule signatures were determined by measuring
m/z 44, 45 and 46 of intact N2O+ ions as well as m/z 30 and 31
of NO+ fragments ions. This allows the determination of average δ15Nav, δ15Nα (δ15N of the
central N position of the N2O molecule) and δ18O (Toyoda and
Yoshida, 1999). δ15Nβ (δ15N of the peripheral N
position of the N2O molecule) is calculated using
δ15Nav= (δ15Nα+δ15Nβ)/2. The 15N site preference
(δ15Nsp) is defined as δ15Nsp= δ15Nα-δ15Nβ. The
scrambling factor and 17O-correction were taken into account (Kaiser and
Röckmann, 2008; Röckmann et al., 2003). Pure N2O (Westfalen,
Münster, Germany) was used as internal reference gas and was analysed in
the laboratory of the Tokyo Institute of Technology using calibration
procedures reported previously (Toyoda and Yoshida, 1999; Westley et al.,
2007). Moreover, the comparison materials from an intercalibration study (S1,
S2) were used to perform a two-point calibration (Mohn et al., 2014). For
correction of non-linear effect due to variable sample amount five different
standard gas mole fractions (0.3, 1, 5, 10, 20 µmol mol-1)
were analysed in each sample run. Samples with similar N2O mole
fractions were run together with at least two standard gases with similar
mole fractions.
All isotopic signatures are expressed as relative deviation (in ‰)
from the 15N / 14N, 17O / 16O and
18O / 16O ratios of the reference materials (i.e. atmospheric
N2 and Vienna Standard Mean Ocean Water (VSMOW), respectively). The
measurement repeatability (1σ) of the internal standard (filled into
vials and measured in the same way as the samples) for measurements of
δ15Nav, δ18O and δ15Nsp
was typically 0.1, 0.1 and 0.5 ‰, respectively.
δ18 O of NO3-
Soil nitrate was extracted in 0.01 M aqueous CaCl2 solution (soil : solution weight
ratio of 1:10) by shaking at room temperature for 1 h.
δ18O of nitrate in the soil solution was determined using the
bacterial denitrification method Casciotti et al., 2002). The measurement
repeatability (1σ) of the international standards (USGS34, USGS35,
IAEA-NO-3) was typically 0.5 ‰ for δ18O.
Δ17O excess in N2O and NO3-
N2O samples collected from soil incubation and N2O produced from
soil NO3- by the bacterial denitrifier method were analysed for
Δ17O using the thermal decomposition method (Kaiser et al., 2007)
with a gold oven (Exp 1.1b, c and 1.2a, b) and with a gold-wire oven
(Exp 1.1a and 2) (Dyckmans et al., 2015). The 17O excess,Δ17O, is defined as (Kaiser et al., 2007)
Δ17O=1+δ17O(1+δ18O)0.5279-1.
The measurement repeatability (1σ) of the international standards
(USGS34, USGS35) was typically 0.5 ‰ for Δ17O.
Soil water analyses
Soil water was extracted with the method described by Königer et
al. (2011) and δ18O of water samples (with respect to VSMOW) was
measured using cavity ring-down spectrometer Picarro L1115-i (Picarro Inc.,
Santa Clara, USA). The measurement repeatability (1σ) of the internal
standards (three calibrated waters with known δ18O: -19.67,
-8.60, +1.37 ‰) was below 0.1 ‰. The overall error
associated with the soil water extraction method determined as standard
deviation (1σ) of the five sample replicates was below 0.5 ‰.
Determination of the extent of isotope exchange
The extent of isotope exchange (x) was determined with two independent
methods described below. In Exp 1 both approaches were applied simultaneously
on the same soil samples, which allowed quantifying the oxygen isotope
exchange with two different methods independently. This enabled the
validation of the 17O excess method, which was used here for the first
time for quantification of isotope exchange. Afterwards this validated method
was applied in the following Exp 2. For both presented methods it is assumed
that after N2O is formed, no further oxygen isotope exchange with
H2O occurs.
δ18O method
This method determines the isotope exchange based on the relative difference
between δ18O of produced N2O and its potential precursors:
soil water and soil nitrate (Snider et al., 2009). To make this method
applicable, parallel incubations with distinct water and/or nitrate isotopic
signatures must be carried out. Therefore, treatments with different water
and nitrate isotopic signatures were applied in Exp 1 (Tables 1, S1). The
calculation is based on two end-member mixing model (water (δw) and nitrate (δn); δ stands for δ18O(N2O)) taking into account the isotope fractionation associated
with O atom incorporation into N2O from each end-member
(εw, fractionation associated with oxygen isotope
exchange with water; εn, fractionation associated with
branching effect during nitrate reduction). This is expressed as
1+δ=x(1+δw)(1+εw)+(1-x)(1+δn)(1+εn)
which can be rearranged to
δ-δn1+δn=x(1+εw)δw-δn1+δn+xεw+(1-x)εn,
where δ-δn1+δn=δ18O(N2O/NO3-) is the dependent variable of the linear regression,
δw-δn1+δn=δ18O(H2O/NO3-) is the independent variable of the linear regression,
x(1+εw) is the slope of the linear regression, approximately equal to
the magnitude of isotope exchange (x), and
xεw+(1-x)εn is the intercept of the
linear regression approximately equal to total fractionation (ε).
Hence, from the linear correlation between δ18O(N2O / NO3-) and δ18O(H2O / NO3-) we can approximate x (the deviation
from the exact value may be up to 0.02, for εw<20 ‰) and the total fractionation ε comprised of both
εw and εn.
Δ17O method
This method determines the isotope exchange based on the comparison of
Δ17O in soil nitrate and produced N2O. It requires the
application of nitrate characterized by high Δ17O. Therefore, soils
were amended with natural NaNO3 Chile saltpeter showing high
Δ17O (ca. 20 ‰) and the Δ17O of the N2O
product was measured. Δ17O of soil water was assumed to be
0 ‰.
The magnitude of oxygen isotope exchange (x) was calculated as
x=1-Δ17O(N2O)Δ17O(NO3-).
The error due to the use of the power-law definition of Δ17O in
combination with a linear mixing relationship (Eq. 4) causes a negligible
relative bias of < 1 % for x.
Correction for N2O reduction
Since δ18O(N2O) values of emitted N2O are strongly
affected by partial N2O reduction, the measured isotope values can only
be informative for the mechanism of N2O production if the reduction is
inhibited or the isotope effects associated with reduction are taken into
account. Exp 1, where we applied both C2H2-inhibited as well as
uninhibited treatments (Table 1), allows us to check the validity of our
correction methods as it directly yields the impact of N2O reduction on
the measured δ18O(N2O) values. In Exp 2, reduction was not
inhibited and the mathematical correction described below was applied.
The correction was made using the Rayleigh fractionation equation
(Mariotti et al., 1981)
1+δS1+δS0=fε,
where δS is the isotopic signature of the remaining substrate (here, measured δ18O of the final, partially reduced, N2O),
δS0 is the initial isotopic signature of the substrate (here,
δ18O of the produced N2O unaffected by the reduction
(δ018O); to be calculated), f is the remaining unreacted fraction (here, the N2O mole fraction f(N2O), directly measured), and
ε is the isotope effect between product and substrate (here,
ε(N2 / N2O), the isotope effect associated with
N2O reduction, taken from the literature; Lewicka-Szczebak et al.,
2014). As it has been shown that the experimental approach largely influences
O isotope effect during reduction (Lewicka-Szczebak et al., 2014, 2015), we
used different ε18O(N2 / N2O) values for static
and flow-through incubations. For the static Exp 1 a mean ε18O(N2 / N2O) value of -17.4 ‰ is used, based on
one common experiment between the study of Lewicka-Szczebak et al. (2014)
(Exp 1) and this study (Exp 1.1). For the flow-through Exp 2 we accept the
ε18O(N2 / N2O) value of -12 ‰ recently
determined for similar flow-through experiments under He / O2
atmosphere (Lewicka-Szczebak et al., 2015). For the correction of
δ15Nsp values one common
ε15Nsp(N2 / N2O) value of
-5 ‰ was used, since it was shown that this value is applicable
for all experimental setups (Lewicka-Szczebak et al., 2014). The error due to
the simplified use of ε15Nsp for the Rayleigh model
(Eq. 5) instead of separate calculations with ε15Nα
and ε15Nβ, causes a negligible bias of the
calculated δ015Nsp values of
< 0.15 ‰ for the presented data set.
N2O isotopic signatures related to water
Relative isotope ratio differences between N2O and soil water,
δ18O(N2O / H2O), were calculated as the difference
between the measured δ18O of produced N2O and of soil water:
δ18O(N2O/H2O)=δ18O(N2O)-δ18O(H2O)1+δ18O(H2O)
In samples where N2O reduction occurred δ18O(N2O / H2O) values were corrected as described above
(Sect. 2.5) and for statistical analyses and modelling exercises the
reduction-corrected values were used (δ018O(N2O / H2O)).
Statistical methods
For results comparisons, an analysis of variance was used with the
significance level α of 0.05. The uncertainty values provided for
the measured parameters represent the standard deviation (1σ) of the
replicates. The propagated uncertainty was calculated using Gauss' error
propagation equation taking into account standard deviations of all
individual parameters.
Results and discussion
Experiment 1
In Table 1 the results are presented as average values from three replicated
incubation vessels with respective standard deviation. Soil nitrate and water
were analysed at the beginning of the experiment from the prepared
homogenized soils, hence no standard deviation but the standard analytical
uncertainty is given.
For different temperature treatments, x(determined by the Δ17O
method) was not significantly different (p= 0.19) but δ018O(N2O / H2O) was slightly higher (p= 0.009) for
8 ∘C ((19.5 ± 0.3) ‰) than for 22 ∘C
((18.6 ± 0.3) ‰) treatment. No significant differences were
observed between the two analysed soil types or between various soil moisture
levels.
When comparing Exp 1.1 and 1.2, x did not show any significant differences,
but the δ018O(N2O / H2O) values were significantly
different (p < 0.001) with higher values for Exp 1.1
((19.1 ± 0.5) ‰) than for Exp 1.2
((16.9 ± 0.8) ‰). It should be noted that the δ18O
values of soil nitrate were much lower in Exp 1.2 (from -2.0 to
6.5 ‰) when compared to Exp 1.1 (from 31.8 to 42.6 ‰) which
might have affected the observed differences in
δ018O(N2O / H2O).
Moreover, for Exp 1 the δ18O method was applied to estimate x
and ε from the relationship between δ18O(N2O / NO3) and δ18O(H2O / NO3)
as described in Sect. 2.4.1.
According to this method, from the linear regression one can decipher x
(slope) and ε (intercept) (Snider et al., 2009). The correlation
is excellent (R2 from 0.989 to 0.997) which indicates that the x and
ε are very stable for all the treatments (Fig. 2). The x is
about 1 (complete exchange) and ε varies from 17.1 (Exp 1.2) to
18.2 ‰ (Exp 1.1). When compared to the results presented in Table 1,
we see slightly higher isotope exchange with the δ18O method when
compared to the Δ17O method. This may be partially due to the fact that
the slope in the δ18O method (Fig. 2) is actually slightly higher than
x (from Eq. (3): x(1+εw)). The difference between
the two experiments is mostly within the error of each method; so far the
results are consistent. The Δ17O method is more useful, since it
allows for individual determinations of x, whereas the correlation obtained
from the δ18O method is based on all data, hence provides a mean
result for x and ε for a whole experiment.
Correlation between oxygen isotopic signatures of N2O and soil
water expressed in relation to soil nitrate, the equation of linear fit
allows for estimation of isotope exchange with soil water (slope of the
linear fit) and the associated isotope effect (intercept of the linear fit).
In red the influence of N2O reduction on the method performance is
presented – red “X” points represent the samples with not inhibited N2O
reduction (note that the slope and intercept are very different), whereas the
red “+” points stand for the same samples after mathematical correction
of N2O reduction effect (as described in Sect. 2.5) which fit very well
to the samples where N2O reduction was inhibited. Data from Exp 1.
Importantly, we found that the δ18O method is not applicable to
samples with uninhibited N2O reduction, if δ18O(N2O)
values are not corrected for N2O reduction. The treatment with
uninhibited reduction of Exp 1.1 was tested and provided very different
results, i.e. largely overestimated x (1.5) and ε (44.8) (red
dashed fit line, Fig. 2). Hence, for proper determination of these factors
the results from treatments with inhibited N2O reduction were used
(solid black fit line, Fig. 2). However, the δ18O values after
mathematical correction for N2O reduction (red “+” points, Fig. 2)
fitted very well to the correlation found for inhibited samples. Hence, the
reduction corrected values (δ018O(N2O)) should rather be
used when applying this method in experiments with uninhibited N2O
reduction. Moreover, in both static experiments we used the C2H2
inhibition technique, and our results indicate almost complete exchange of
oxygen isotopes with soil water, which indicates that the isotope exchange
process is not inhibited by C2H2 addition.
Experiment 2
In Table 2 the results are presented as average values from three replicate
incubation vessels with respective standard deviation. The extent of oxygen
isotope exchange (x) ranges from 55 to 85 % and is lower and much more
variable when compared to Exp 1. δ018O(N2O / H2O)
varies between 18.6 and 36.9 ‰, which is significantly higher when
compared to the values determined in Exp 1.
Oxygen isotope effects at nearly complete isotope exchange
In the case of very high, almost complete, isotope exchange with soil water
(Exp 1), the relative isotope ratio difference between N2O and H2O
(δ018O(N2O / H2O)) is quite stable and ranges from
15.6 to 19.8 ‰ (Table 1). In contrast, the relative isotope ratio
difference between N2O and NO3- (δ018O(N2O / NO3-)) shows large variations from -36.1
to 18.0 ‰ (Fig. 3).
Relation between relative isotope ratio differences between produced
N2O and soil water (δ018O(N2O / H2O) and
between produced N2O and soil nitrate (δ018O(N2O / NO3-); on the right δ18O values
of the initial soil nitrate for different treatments. δ18O values
of the initial soil water ranged between -13.5 and -1.6 ‰ (see
Table 1) and its variation had no impact on
δ018O(N2O / H2O). Open symbols: treatments with
synthetic nitrate as fertilizer, filled symbols: treatments with natural
Chile saltpeter as fertilizer. Data from Exp 1.
The ε determined in Fig. 2 represents theoretically the total oxygen
isotope fractionation (from Eq. (3): xεw+ (1-x)εn), but in the case of the nearly whole isotope
exchange (x=1) ε equals εw and
εw = (δN2O-δw) / (δw+1)
=δ18O(N2O / H2O), hence both the intercept in
Fig. 2 and δ18O(N2O / H2O) in Fig. 3 should provide
rough estimates for εw. However, for x< 1
δ18O(N2O / H2O) depends also on δn
and εn and the intercept (Fig. 2) includes εn. Both these values indicate a slight difference between both
experiments: for Exp 1.1 ε of (18.2 ± 0.6) (intercept,
Fig. 2) and δ18O(N2O / H2O) of (19.1 ± 0.5)
(mean ± SD, Table 1) are higher than for Exp 1.2, (17.1 ± 0.3)
and (16.7 ± 0.8), respectively. This slight difference is most probably
due to x slightly lower than 1, as indicated by Δ17O method and
additional impact of δn and εn. It can be
noted that δ018O(N2O / H2O) slightly increases
with higher δ18O values of nitrate (Fig. 3), i.e. the difference of
about 40 ‰ in δ18O of applied NO3- results in about
2 ‰ change in δ018O(N2O / H2O). Hence,
only about 5 % of the difference in nitrate isotopic signature is
reflected in the produced N2O, suggesting that an equivalent percentage
of O(N2O) originated from NO3-. This is very consistent with the
determined extent of isotope exchange with soil water, which was
(95.6 ± 2.6) % (Table 1).
Taken together, the data indicate that the δ18O(N2O)
values are clearly influenced by the δ18O of soil water,
whereas δ18O of soil nitrates has only very little influence.
Hence, the O isotope fractionation during N2O production by
denitrification should be considered in relation to soil water, rather than
soil nitrates.
Oxygen isotope effects at variable isotope exchange
In contrast to Sect. 3.3, x was more variable for the flow-through
incubation (Exp 2) and also significantly lower. In general, lower x was
associated with higher δ018O(N2O / H2O) values. In
Fig. 4 we can compare results from static incubations (red symbols) with the
flow-through incubations (black symbols). This comparison clearly shows that
the pattern of isotope exchange and associated oxygen fractionation differs
significantly between both experimental approaches. The essential difference
in Exp 2 was the use of a flow-through system with an oxic atmosphere at the
beginning of the incubation (though results presented originate from the
anoxic phase). This resulted in lower production rates for N2O when
comparing the respective soil (Tables 1 and 2), e.g.
80 µg kg-1 h-1 (mass of N as sum of N2O and
N2 per mass of dry soil) for the silt loam soil at 80 % WFPS in
Exp 2.3 but 261 µg kg-1 h-1 in Exp 1.1c. This may
suggest an impact of N2O production rate on extent of isotope exchange.
However, for static anoxic incubations the effect of production rate was not
observed, e.g. between Exp 1.1a and 1.1b (Table 1), where we have different
production rates but similar x and δ018O(N2O / H2O).
δ018O(N2O / H2O) as a function of isotope
exchange extent, x (determined with Δ17O method). Red symbols:
Exp 1, black symbols: Exp 2; open symbols: incubations with lower WFPS
(70 %), filled symbols: incubations with higher WFPS (80 %). Note
that same symbol shapes always represent the same soil.
Interestingly, the correlation between x and δ018O(N2O / H2O) seems to differ for different soil
types. Very clearly both sandy soils represent distinct and weaker
correlation when compared to silt loam and organic soil. Most probably this
is due to different oxygen fractionation pattern during N2O formation in
both soils, which we try to elucidate in the theoretical model presented
below.
The mechanism of oxygen isotope fractionation – a fractionation
model
To better understand the mechanism of oxygen isotope fractionation and the
relation between the apparent isotope effect and the extent of isotope
exchange we applied a simulation calculation where the total isotope effect
was calculated from the theoretical isotope fractionation associated with two
enzymatic reduction steps: NIR and NOR. This model was based on the
calculations presented by Rohe et al. (2014a) for pure fungal cultures, where
this approach has been described in detail. The model assumes that
δ18O(N2O) is determined by two isotope fractionation processes
associated (i) with the branching isotope effect (εn)
and (ii) with the isotope effect due to isotope exchange with soil water
(εw), both possible at NIR or NOR. This can be expressed
by the following isotope mass balance equations:
1+δ=xNOR(1+δw)(1+εw)+(1-xNOR)(1+δNO)(1+εNOR)1+δNO=xNIR(1+δw)(1+εw)+(1-xNIR)(1+δn)(1+εNIR),
where
1-x=(1-xNIR)(1-xNOR)1+εn=(1+εNIR)(1+εNOR).
After substitution and transformation, this gives
δ-δw1+δw=(1-x)(1+εn)δn-δw1+δw+(x-xNOR)εNOR(1+εw)+xεw+(1-x)εn.
We neglected the possible fractionation associated with the NAR reduction,
i.e. δ(NO2-)=δ(NO3-)=δn in Eq. (11).
This enzymatic step was investigated by Rohe et al. (2014a), and appeared to
have no significant impact on the total oxygen fractionation, i.e. the
branching fractionation for nitrate treatments was in no case higher than for
nitrite treatment. This indicates that the oxygen fractionation between
nitrate and nitrite is low due to cancellation of the intramolecular effect
of about 30 ‰ (Casciotti et al., 2007) by the intermolecular effect
when the nitrate pool is not completely consumed. Hence, we only focused here
on differentiating between NIR and NOR enzymatic reduction steps, which are
most likely the enzymatic reactions crucial for determining final N2O
isotopic values (Kool et al., 2007).
There are many unknown factors in Eq. (11); first of all, isotopic
fractionation factors εn and εw. We
have compiled the results of both methods applied for Exp 1 data – the δ18O method and the Δ17O method – to estimate these factors. Using the
δ18O method, ε was determined from the intercept in
Fig. 2 and this value represents total fractionation: ε=xεw+ (1-x) εn (see Sect. 2.4.1).
Using the Δ17O method, individual x values were calculated for
each sample. We have also measured δ18O(N2O / H2O)
and δ18O(NO3- / H2O) for each sample, hence from
the transformed Eq. (3):
δ-δw1+δw=(1-x)(1+εn)δn-δw1+δw+xεw+(1-x)εn
and knowing that xεw+ (1-x) εn= 0.0182 for Exp 1.1 and xεw+ (1-x)
εn = 0.0171 for Exp 1.2 (Fig. 2) we have calculated
εw and εn for each sample. Table 3
summarizes the results.
Isotopic fractionation factors calculated based on Exp 1 results
with Eq. (12) (see text for details). Results presented separately for
Exp 1.1 and 1.2 and mean values for both.
εw (‰)
εn (‰)
Exp 1.1
17.44 ± 0.71
0.74 ± 0.70
Exp 1.2
17.50 ± 0.67
-0.39 ± 0.66
mean all
17.48 ± 0.66
0.03 ± 0.86
The determination of εw is very precise, with no
significant difference between Exp 1.1 and 1.2 (p=0.868). The value
obtained (17.5 ± 0.7) ‰ is within the range of the previous
values determined for chemical exchange ε(NO2- / H2O) = 14 ‰ and ε(NO3- / H2O) = 23 ‰ (Böhlke et al., 2003;
Casciotti et al., 2007). So far there are no data for the isotope effect of
chemical exchange ε(NO / H2O). Therefore, we assumed equal
εw values for isotope exchange associated with NIR and
NOR, similarly to previous studies (Rohe et al., 2014a; Snider et al., 2012).
Hence, the εw value determined here is a hypothetical
mean value of enzymatically mediated isotope exchange associated with NIR
(εw(NO2- / H2O)) and NOR
(εw(NO / H2O)).
The εn is also quite stable with a weak (p=0.006) and very
small (below 1 ‰) difference between Exp 1.1 and 1.2. The
εn values found are very low and vary around 0, from -1.9
to 2.1 ‰. This is much lower than in previous studies, which
reported εn from 10 to 30 ‰ (Casciotti et al.,
2007; Rohe et al., 2014a).
We checked how well these calculated values fit for the individual samples of
both experiments. We started with the simplest Scenario 0, where we assume
the values determined in Table 3 for εw and
εn and calculate the δ18O(N2O) with
Eq. (11), which is then compared with the measured δ18O(N2O)
and the difference between measured and calculated δ18O(N2O)
value (D) is determined (Table 4). Since the mean value of 0 was assumed
for εn in this scenario, the isotope exchange can be
associated either with NIR or NOR without any effect on the final
δ18O(N2O), because Eq. (11) is simplified to
δ-δw1+δw=(1-x)δn-δw1+δw+xεw.
This scenario works quite well for Exp 1 data with the maximal D of
1.4 ‰. However, for Exp 2 data we obtain significant overestimation
of the calculated δ18O(N2O) values for sandy soils (Exp 2.1
and 2.2) up to 6.1 ‰ and underestimation for two other soils,
reaching up to 12.2 ‰ for organic soil (Exp 2.5). Why does the model
developed based on Exp 1 data not work for Exp 2 data? We expect that the
εw value should be quite stable for all the samples. It
was observed in the study by Casciotti et al. (2007) that ε(NO2- / H2O) values varied in a very narrow range. Also in
our study in Fig. 2 we obtained very good correlation with stable slope which
suggests that the εw value must be very stable and almost
identical for all the samples. It can be supposed that
εn values can be more variable, but due to nearly
complete isotope exchange in Exp 1 these potential variations cannot be
reflected in δ18O(N2O) values. Also, the study by Rohe et
al. (2014a) indicated possibly wide variations of εn from
10 to 30 ‰.
Oxygen fractionation model based on the results obtained (δ018O(N2O)) and isotope exchange (x) determined by Δ17O method and εw= 17.5 ‰ determined from
Exp 1 data (Table 3). Scenarios with varied εn values
and xNIR or xNOR (fraction of isotope exchange associated
with NIR or NOR) are compared. D is the difference between measured
δ18O of N2O and the calculated δ18O of N2O in a
particular scenario.
Scenario 0
Scenario 1
Scenario 2
Scenario 3
x=xNIR or
xNIR=x;
xNIR=0;
xNIR=xNOR
xNOR
xNOR=0
xNOR=x
εn=0
εn fitted
εn fitted
εn fitted
εw=17.5 (‰)
εw=17.5 (‰)
εw=17.5 (‰)
εw=17.5 (‰)
D
εn
D
εn
D
εn
D
Exp 1.1a
0.2
0.3
0.00
2.3
0.00
1.0
0.00
0.6
1.2
0.00
16.0
0.00
5.3
0.00
Exp 1.1b
0.1
0.2
0.00
2.7
0.00
0.9
0.00
-1.2
-2.3
0.00
-22.6
0.00
-8.6
0.00
Exp 1.1c
0.2
0.4
0.00
4.7
0.00
1.7
0.00
0.0
0.1
0.00
0.6
0.00
0.2
0.00
Exp 1.2a
-0.3
-0.5
0.00
-3.7
0.00
-1.6
0.00
-0.8
-1.5
0.00
-18.4
0.00
-6.2
0.00
0.3
0.6
0.00
4.5
0.00
1.9
0.00
0.2
0.3
0.00
2.7
0.00
1.0
0.00
Exp 1.2b
-0.4
-0.7
0.00
-4.0
0.00
-1.9
0.00
0.1
0.2
0.00
1.7
0.00
0.7
0.00
1.4
2.6
0.00
38.5
0.00
12.1
0.00
-0.7
-1.3
0.00
-72.8
0.00
-12.5
0.00
-0.3
-0.6
0.00
-19.3
0.00
-4.2
0.00
0.2
0.4
0.00
0.0
0.22
0.0
0.22
Exp 2.1
-4.0
-6.2
0.00
-14.7
0.00
-10.0
0.00
-5.3
-8.2
0.00
-19.9
0.00
-13.4
0.00
Exp 2.2
-5.2
-7.6
0.00
-15.0
0.00
-11.0
0.00
-6.1
-9.1
0.00
-20.0
0.00
-14.1
0.00
Exp 2.3
2.5
3.2
0.00
4.9
0.00
4.0
0.00
3.0
3.8
0.00
5.7
0.00
4.7
0.00
Exp 2.4
1.1
1.6
0.00
3.4
0.00
2.4
0.00
2.2
2.9
0.00
4.8
0.00
3.8
0.00
Exp 2.5
2.8
4.2
0.00
8.5
0.00
6.2
0.00
12.2
18.1
0.00
37.1
0.00
27.0
0.00
Exp 2.6
2.2
3.8
0.00
20.9
0.00
10.2
0.00
4.2
6.8
0.00
19.1
0.00
12.2
0.00
Therefore, for the next scenarios (Scenario 1, 2 and 3 – Table 4) we assumed a
stable εw value of 17.5 ‰, as determined from
Exp 1 (Table 3), and εn values were calculated
individually for each sample with Eq. (11) from the δ018O(N2O / H2O) values. In each scenario εn was equally distributed between NIR and NOR according to Eq. (10),
so that εNIR=εNOR. For our samples we
know the value of total isotope exchange (x determined with Δ17O method), but we do not know at which enzymatic step(s) this exchange
occurred. Since the isotope exchange has very different impact on the final
δ18O(N2O) when associated with NIR or NOR, we can obtain this
information by comparing different scenarios (Table 4). In Scenario 1 the
total isotope exchange is associated with the first reduction step NIR and in
Scenario 2, with the final reduction step NOR. In Scenario 3 the total
isotope exchange is equally distributed between both steps NIR and NOR
according to Eq. (9) so that xNIR=xNOR.
In this study, we could not determine at which enzymatic step isotope
exchange occurs, but only its impact on the implied isotope effects. Namely,
in Scenario 1 the exchange effect associated with xNIR precedes the
branching effect at NOR (εNOR) and, conversely, in
Scenario 2 the exchange isotope effect associated with xNOR occurs
after both branching effects (εNIR,
εNOR). Hence, in Scenario 1 εNOR has
a more direct impact on the final δ18O(N2O) whereas in
Scenario 2 the last fractionation step is related to εw
(Eq. 11). Therefore, applying different scenarios results in different values
for the calculated εn (Table 4).
The narrowest range of variations of the calculated εn
values was obtained in Scenario 1. For Exp 1 they vary around 0, similarly to
the results presented in Table 3, which indicates that this model and the
equations applied for δ18O method (Eq. 12) are actually the same.
For Exp 2 the calculated εn values are negative for sandy
soils (Exp 2.1 and 2.2) from -9.1 to -6.2 ‰ and positive for
other soils with lower values for silt loam from 1.6 to 3.8 ‰ and
higher for organic soil from 3.8 to 18.1 ‰ (Table 4). Variations of
calculated εn values are much larger in Scenario 2 with
a particularly wide range for Exp 1 from -72.8 to +38.5 ‰. For
Exp 2, a similar trend as in Scenario 1 is observed, with negative values for
sandy soils (down to -20.0 ‰) and highest values for organic soil
(up to 37.1 ‰). The absolute values are generally larger and the
variations among them are thereby increased when compared to Scenario 1. The
strongly negative εn values obtained for Scenario 2 are
outside the plausible range based on previous determinations
(Casciotti et al., 2007; Rohe et al., 2014a). Moreover, for the last sample
of Exp 1 where x=1 this scenario fails in finding the
εn value for D=0, because for complete isotope exchange
by NOR, the associated branching isotope effect has no impact on the final
δ18O(N2O). Hence, Scenario 1 is more plausible because
(i) the overall εn variations are smaller and (ii) we do
not find extremely negative values. Results from Scenario 3 are situated in
the middle of Scenario 1 and 2, and show larger variations than Scenario 1,
but without the extreme outliers, hence it can be also a plausible model. From
comparison of these scenarios we can say that isotope exchange is likely
associated with NIR and may also partially take place at NOR (but not NOR
alone). This reinforces the previous findings from pure culture studies which
suggested that the majority of isotope exchange is associated mainly with nitrite
reduction (Garber and Hollocher, 1982; Rohe et al., 2014a). Moreover, each
scenario indicates clearly a much lower branching effect for the two sandy
soils in Exp 2 when compared to silt loam and organic soil. This is the
reason behind the different slope of correlation δ018O(N2O / H2O) vs. x in Fig. 4 for sandy soils.
Lower εn values mean that N2O is less enriched in
18O in relation to soil nitrate and lower x results in smaller
increase in δ18O(N2O) values, which was observed for sandy
soils (Fig. 4).
For each scenario our model indicated rather lower εn
values than previously assumed (Casciotti et al., 2007; Rohe et al., 2014a).
But actually, the isotope effect determined by Casciotti et al. (2007), +25
to +30 ‰, takes only the intramolecular branching effect into
account, because in the bacterial denitrification method the whole nitrate
pool is quantitatively consumed, hence the intermolecular isotope effect
cannot manifest. Therefore, the values found by Casciotti et al. (2007)
represent the maximal possible branching effect. In the experiment presented
by Rohe et al. (2014a) only very little added substrate was reduced, hence we
should also observe the intermolecular isotope effects. Indeed, the model
applied by Rohe et al. (2014a) indicated lower magnitudes for net branching,
down to +10 ‰ for εNIR and 0 ‰ for
εNAR. This may suggest that the net branching effect
decreases with smaller reaction rates because of intermolecular isotope
effects. But are negative net branching effects actually possible? The answer
is yes, provided that the intermolecular effect exceeds the intramolecular
effect, i.e. the former must be more negative than -30 ‰. An idea
about the magnitude of the intermolecular effect can be obtained from the
change in isotopic signature of the remaining nitrate, since this reflects
the enrichment in residual nitrate-18O due to intermolecular effects. In
pure culture studies this effect ranges from -23 to -5 ‰ (Granger
et al., 2008), but in soil incubations values as low as -37 ‰ have
been observed (Exp. 1F in Lewicka-Szczebak et al., 2014). Hence, slightly
negative net εn values are theoretically possible, but up
to a few ‰ for each enzymatic step, which gives the minimal overall
εn of about -10 ‰. Therefore, the results of
Scenario 2 must be rejected, whereas the values found in Scenario 1 are most
plausible.
Significance for quantification and differentiation of soil
denitrification
From the presented results it is most surprising and incomprehensible why
the same soils show various extents of isotope exchange with soil water, and
especially, why this exchange was high and stable under static anoxic
conditions and significantly lower in flow-through incubations. Most
probably, in the static inhibited experiments denitrification is the only
N2O producing process and in the flow-through uninhibited incubations
other N2O producing processes may significantly contribute to N2O
production. These incubations were performed initially under oxic conditions,
which were switched to anoxic conditions after 3 days. However, all the
results presented here originate from this anoxic phase, since the N2O
production during oxic phase was too low for Δ17O analyses. Hence,
the potentially contributing processes might be fungal denitrification,
co-denitrification, nitrifier denitrification or dissimilatory nitrate
reduction to ammonium (DNRA). 15N site preference
(δ15Nsp) may be used as a tracer to distinguish some of
these processes. It is known that fungal denitrification and nitrification
are characterized by significantly higher δ15Nsp values
(33 to 37 ‰, Rohe et al., 2014a; Sutka et al., 2006, 2008) when
compared to bacterial denitrification and nitrifier denitrification (-11 to
0 ‰, Sutka et al., 2006; Toyoda et al., 2005). To check the
hypothesis of mixing of N2O from various sources we plotted
δ018O (N2O / H2O) values against
δ015Nsp values of produced N2O (Fig. 5).
Relation between δ015Nsp of produced N2O
and relative ratio difference between produced N2O and soil water
(δ018O(N2O / H2O)). Red symbols: Exp 1; black
symbols: Exp 2; open symbols: incubations with lower WFPS (70 %); filled
symbols: incubations with higher WFPS (80 %). Note that the same symbol
shapes always represent the same soil. Grey dashed lines represent the
possible range of linear fit when extreme values of isotope effects for
N2O reduction are assumed in correction calculations (Eq. 5). Range of
values for fungal denitrification from Rohe et al. (2014a).
It can be clearly noticed that the results from the inhibited experiment
(Exp 1, red symbols) fit perfectly into the field of bacterial
denitrification. Similarly, the results of sandy soils from Exp 2 show a
slightly wider range, but still are typical for bacterial denitrification. In
contrast, silt loam soil (Exp 2.3, 2.4) and organic soil (Exp 2.5, 2.6) both
show increased δ018O(N2O / H2O) and δ015Nsp values which are very well correlated. This could
indicate that in Exp 2 another process characterized by high δ15Nsp and δ18O values has significant contribution to
total N2O production by these two soils. This could be nitrification,
which is not very plausible due to the anoxic conditions, or fungal
denitrification. But it remains unclear why this was not observed in the
inhibited static incubation for the same soil (silt loam). C2H2
inhibition does not affect fungal denitrification (Maeda et al., 2015) in that NO3- and NO2- availability is not restricted by inhibited
nitrification. However, in the flow-through incubations, the first oxic phase
might have activated other microorganisms, possibly preferentially fungi.
This could explain why their contribution is observed in Exp 2 but not
in Exp 1. Such an activation of denitrification by oxygen supply has been
documented for one fungus species (Zhou et al., 2001).
We checked whether the correlation presented in Fig. 5 could have resulted from
calculation artifacts, since all of the higher δ018O(N2O / H2O) and δ015Nsp
values were corrected for N2O reduction (according to the method
described in Sect. 2.5). This correction method does not provide very precise
results, since the isotope effects associated with N2O reduction are not
entirely stable and predictable (Lewicka-Szczebak et al., 2014, 2015).
Therefore, we checked whether this correlation may be only a calculation
artifact and recalculated the values assuming larger range of isotopic
fractionations (±5 ‰, resulting in ε15Nsp(N2 / N2O) from -10 to 0 ‰ and
ε18O(N2 / N2O) from -20 to -6 ‰).
Results show that the correlation may slightly change in slope (from 0.41 to
0.85), intercept (from -10.4 to -18.0) and significance (R2 from
0.64 to 0.91). But it always keeps the same trend, i.e. for Exps 2.3–2.6
we obtain in any case a correlated increase of δ015Nsp
and δ018O(N2O / H2O) (see grey dashed lines in
Fig. 5), proving that the indication for further contributing processes
cannot be an artifact of the correction approach. For these experiments
(2.3–2.6) in our model calculations (Table 4) higher
εn values were always found when compared to Exp 1 and 2.1–2.2.
For pure culture studies of fungal denitrification the
εn values determined by a similar modelling approach were also
higher, up to 30 ‰ (Rohe et al., 2014a). This would support the
hypothesis on fungal denitrification contribution.
Source of Δ17O in atmospheric N2O
In Exp 1 the Δ17O(N2O) values obtained from all measured
N2O samples were very low. Moreover, we also included the treatment with
chemical nitrate as fertilizer, characterized by slightly negative
Δ17O excess (of ca. -1.5 ‰), and the produced N2O did
not show any positive Δ17O excess (results not shown). The produced
N2O is always characterized by smaller 17O excess (Δ17O
values closer to 0) than in the source nitrate (Table 1). These results
indicate that denitrification produces N2O of randomly distributed
oxygen, due to a mostly very high extent of isotope exchange with soil water
and the consequent loss of 17O excess of nitrate. However, in Exp 2
numerous samples showed lower extent of isotope exchange, down to 50 %,
and the 17O excess of nitrate is partially transferred to N2O,
resulting in Δ17O(N2O) up to 5 ‰. This indicates
that denitrification may be the source of atmospheric N2O
with 17O excess, as previously supposed (Kaiser et al., 2004; Michalski
et al., 2003), but the magnitude of this excess is largely reduced by the
exchange of oxygen isotopes with randomly distributed soil water.