Constraining the oxygen isotopic composition of nitrate produced by

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Constraining the oxygen isotopic composition of nitrate produced by nitrification Danielle Boshers, Julie Granger, Craig Tobias, John Karl Bohlke, and Richard L Smith Environ. Sci. Technol., Just Accepted Manuscript • DOI: 10.1021/acs.est.8b03386 • Publication Date (Web): 03 Jan 2019 Downloaded from http://pubs.acs.org on January 4, 2019

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Environmental Science & Technology

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Constraining the oxygen isotopic composition of nitrate produced by nitrification

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Danielle S. Boshers*†, Julie Granger†, Craig R. Tobias†, John K. Böhlke§, and Richard L. Smith ‡

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†Department

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§

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of Marine Sciences, University of Connecticut, 1080 Shennecossett Road, Groton, Connecticut 06340 U.S. Geological Survey, 431 National Center, Reston, Virginia 20192

‡U.S.

Geological Survey, 3215 Marine Street, Boulder, Colorado 80303

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Corresponding Author

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*e-mail: [email protected]; phone: (269) 719-0456

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Abstract

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Measurements of the stable isotope ratios of nitrogen (15N/14N) and oxygen (18O/16O) in nitrate

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(NO3-) enable identification of sources, dispersal, and fate of natural and contaminant NO3- in

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aquatic environments. The 18O/16O of NO3- produced by nitrification is often assumed to reflect

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the proportional contribution of oxygen atom sources, water and molecular oxygen, in a 2:1 ratio.

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Culture and seawater incubations, however, indicate oxygen isotopic equilibration between

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nitrite (NO2-) and water, and kinetic isotope effects for oxygen atom incorporation, which

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modulate the NO3- 18O/16O produced during nitrification. To investigate the influence of kinetic

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and equilibrium effects on the isotopic composition of NO3- produced from the nitrification of

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ammonia (NH3), we incubated stream water supplemented with ammonium (NH4+) and

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increments of 18O-enriched water. Resulting NO3- 18O/16O ratios showed (1) a disproportionate 1

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sensitivity to the 18O/16O ratio of water, mediated by isotopic equilibration between water and

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NO2-, as well as (2) kinetic isotope discrimination during O-atom incorporation from molecular

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oxygen and water. Empirically, the NO3- 18O/16O ratios thus produced fortuitously converge near

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the 18O/16O ratio of water. More elevated NO3- 18O/16O values commonly reported in soils and

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oxic groundwater may thus derive from processes additional to nitrification, including NO3-

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reduction.

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Introduction

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Human activity has greatly altered the nitrogen cycle through industrial nitrogen fertilizer

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use, fossil fuel combustion, and wastewater disposal, contributing to the deterioration of

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groundwater quality, and to the eutrophication of lakes, rivers, estuaries, and shelf seas. In order

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to elucidate specific causes of water quality impairment and eutrophication, the sources and fate

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of reactive nitrogen in the environment need to be understood. To this end, the naturally

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occurring stable isotope ratios of nitrogen (15N/14N) and oxygen (18O/16O) in nitrate (NO3-),

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henceforth expressed in “delta” notation as 15N vs. air and 18O vs. V-SMOW1, provide a useful

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tracer to determine the origins, dispersal, and biogeochemical transformations of natural and

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contaminant NO3-. To use NO3- isotopes in this way, it is necessary to know the isotopic

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composition of source end-members, as well as the extent to which NO3- isotopologues are

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fractionated by specific biological reactions that produce and consume NO3-. Isotopic

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fractionation during an irreversible reaction is described by the kinetic isotope effect, , where

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*k (‰) = (k/*k -1)1000, and k and k* are the respective reaction rate constants of light and

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heavy isotopologues.1

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NO3- is produced in the environment by nitrification, the biological oxidation of ammonia

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(NH3) to nitrite (NO2-) then NO3-, by chemoautotrophic bacteria and archaea.2,3 NO3- produced by

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nitrification of natural and contaminant NH3 is distinguished from other NO3- sources, namely

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atmospheric and industrial fertilizer NO3-, by its comparatively low 18ONO3 values.4 The latter are

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assumed to derive from the respective O atoms originating from molecular O2 (approximately 24

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‰ vs. V-SMOW in air and air-saturated water) and H2O (typically -15 ‰ to 0 ‰ vs. V-SMOW in

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tropical-temperate meteoric water and the ocean) 5,6,7 during the biological oxidation of NH3 to

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NO3-. Oxygen isotopic tracer studies of nitrification indicate that approximately one O atom each

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from O2 and H2O contribute to formation of NO2- from NH3 (NO2--nit) , and that one O atom from

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H2O is added during further oxidation of NO2- to NO3- (NO3--nit ).8,9 These results have provided

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a commonly used benchmark for estimating 18O values of NO2- and NO3- produced by

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nitrification in the environment10:

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𝛿18𝑂𝑁𝑂2 ― 𝑛𝑖𝑡 =

1 2

1

(1)

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𝛿18𝑂𝑁𝑂3 ― 𝑛𝑖𝑡 =

2 3

1

(2)

(𝛿18𝑂𝐻2𝑂) + 2(𝛿18𝑂𝑂2) (𝛿18𝑂𝐻2𝑂) + 3(𝛿18𝑂𝑂2)

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By this convention, the 18O of NO3- produced by nitrification in systems where the 18O of

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water ranges from -15 ‰ to 0 ‰ should range between about -2 ‰ and 8 ‰, assuming a 18O

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value of ~24 ‰ for molecular O2. Studies of NO3- in soils, groundwaters, and streams seem to

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indicate that the 18O of terrestrial NO3--nit may conform to the model described by Eq. 2 in some

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cases but not in others. Instances where the 18O of NO3--nit is higher than predicted by Eq. 2

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have been attributed to mechanisms that cause O isotope enrichments of H2O and O2, such as

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evaporation and respiration.11–13 Values lower than predicted by Eq. 2 have been hypothesized

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to derive from isotopic exchange of O atoms between the NO2- intermediate and H2O and to

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kinetic O isotope fractionations during nitrification.14,15 An additional complication is that

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reported NO3- 18O values are not all comparable because some were determined before reliable

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calibration schemes were available, in some cases resulting in anomalously high apparent 18O

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values for NO3--nit.16,17

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The variability among reported isotopic dynamics of nitrification reflects a lack of

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fundamental consensus regarding the processes that result in 18O variations of NO3- produced

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by nitrification, leading to uncertainty in interpretations of 18O signatures of NO3- in the 4


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Figure 1. Schematic diagram of O isotopic fractionation and exchange during nitrification. Sources of O atoms (O2 and H2O) for NH3 and NO2- oxidation are shown, as well as putative isotope effects associated with oxygen atom incorporation (18k,O2, 18k,H2O,1, and 18k,H2O,2), NO2- equilibration (18eq) and NO2- oxidation (18k,NO2). Also shown is the fraction of NO2- O atoms that have equilibrated with H2O, XNO2.

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environment. Recent experimental studies with isolated nitrifier cultures and seawater

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incubations reveal that 18ONO3 values are sensitive not only to the O isotopic composition of the

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H2O and O2 incorporated during nitrification, but also to kinetic and equilibrium isotope effects

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during the reactions (Fig. 1). The respective incorporations of O atoms from molecular O2 and

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H2O during NH3 oxidation by cultures of bacterial and archaeal isolates from marine and

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terrestrial systems are associated with substantial O isotope fractionations (18k,H2O,1 and

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18

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oxidation is associated with a temperature-dependent equilibrium isotope effect of ~13 ‰ (Fig.

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1).18,20 During NO2- oxidation to NO3-, the O atoms of NO2- are subject to an inverse kinetic isotope

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effect (18k,NO2)21, and the concurrent O atom incorporation from water also involves isotopic

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discrimination (18k,H2O,2; Fig. 1).19,21 Although the isotope effects characterized in cultures and in

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incubations of marine assemblages are sizeable, they are seldom considered when inferring the

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18ONO3 associated with nitrification in terrestrial aquatic systems. Thus, interpretations of NO3-

k,O2).

18,19

Abiotic or biologically-enhanced O atom exchange between NO2- and H2O during NH3

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isotope distributions in the environment are often incongruous with observations from cultures

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and incubations.

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Accurate interpretations of NO3- isotope distributions in the environment require a sound

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mechanistic understanding of factors that modulate the isotopic composition of NO3- produced

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by nitrification. Here, we investigate the influence of the 18O of ambient H2O and of O isotope

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fractionations on the isotopic composition of NO3- produced during nitrification of NH3 by natural

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freshwater and brackish microbial consortia in order to determine whether the 18ONO3 produced

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is adequately described by the weighted sum of 18O source contributions (Eq. 2) or if kinetic

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isotope effects and/or NO2- isotopic equilibration also need to be considered. In doing so, we aim

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to reconcile the apparent incongruity between interpretations of the 18ONO3 produced by

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nitrification in pure cultures and marine assemblages compared to common observations in

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freshwater aquatic systems, and to provide a robust framework from which to distinguish

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nitrified NO3- in the environment from other NO3- sources.

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Materials and Methods

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Stream water incubations

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Water containing ambient microbial communities was obtained from two streams in coastal

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Connecticut in 2016: an upland freshwater stream in the spring in a preliminary experiment

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(Experiment 1), results from which are largely presented in Supplementary Materials, then from

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a brackish tidal stream (salinity of 18 g/kg) in the fall (Experiment 2). Unfiltered water was

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collected in acid-washed polypropylene carboys, then dispensed through a multi-layered coarse

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mesh into twelve acid-washed, sterile 2 L glass media bottles. Each bottle was amended with

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NH4Cl to a concentration of 50 mol L-1. The 18OH2O of each of 4 treatments was adjusted in

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triplicate bottles with incremental additions of 97-atom-% 18O-labeled water (Cambridge Isotope

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Laboratories, Lot #: WP-14-25; Table 1). Bottles were capped loosely to permit air exchange and

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incubated statically in the dark at room temperature. Incubations were mixed and subsampled

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weekly until the appearance of NO2-, and then quasi daily in Experiment 2 until complete

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conversion of NH4+ to NO3-, but only weekly in Experiment 1. Given the slow growth rates of

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nitrifiers in our incubations (see Results), the continuous exchange of headspace with air, and

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episodic agitation of the incubations during samplings, the 18O of dissolved O2 was assumed to

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remain close to the air-equilibrium value throughout the experiments. Samples for nutrient

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concentration and NO3- isotope measurements were frozen immediately upon collection pending

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analysis. Samples for NO2- isotope analysis were analyzed immediately in Experiment 1, or

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adjusted to a pH of ~10 with 1 M sodium hydroxide and stored frozen pending analysis in

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Experiment 2.

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Dissolved inorganic nitrogen (DIN) analyses.

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Nutrient concentrations (NO3-, NO2-, NH4+) were analyzed at the University of Connecticut

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(UConn) on a SmartChem® nutrient analyser, (Unity Scientific, Brookfield, Connecticut), using

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standard spectrophotometric protocols.22–25 Some NO2- and NO3- samples were also analyzed by

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chemiluminescent detection on a NOx analyzer (model T200 Teledyne Advanced Pollution

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Instrument) following reduction to nitric oxide (NO) in a heated iodine solution for NO2- and

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reduction in a heated vanadium (III) solution for NO3- plus NO2-.26,27

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Table 1. 18OH2O of experimental treatments (mean ±  of triplicate bottles).

Treatment 1 2 3 4 129

Experiment 1 18OH2O (‰) -6.9 ± 0.0 -6.6 ± 0.0 -5.5 ± 0.0 -4.2 ± 0.1

Experiment 2 18OH2O (‰) -2.2 ± 0.0 3.8 ± 0.2 14.6 ± 0.2 34.0 ± 0.3

Isotope ratio analyses

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NO3- N and O isotope ratios (15NNO3 vs. air and 18ONO3 vs. V-SMOW) were measured at

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UConn by conversion to N2O with the denitrifier method.28–30 N2O was analyzed on a Thermo

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Delta V gas chromatograph isotope ratio mass spectrometer (GC-IRMS) interfaced with a custom-

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modified Gas Bench II and sample preparation device, dual cold traps, and GC Pal auto-sampler.

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Prior to analysis of NO3-, NO2- was removed from samples with sulfamic acid where [NO2-] ≤ [NO3-

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], or with ascorbic acid where [NO2-] > [NO3-], as the latter results in more accurate 15NNO3 values

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at relatively elevated [NO2-] relative to [NO3-].31,32 Sample analyses were calibrated using NO3-

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reference materials IAEA-N3 and USGS34.17,33 Sample volumes and concentrations were matched

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to those of standards to minimize potential matrix effects on 18ONO3.34 The 18O of water among

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treatments was taken into account to correct for isotopic exchange during conversion to N2O.

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Replicate batch analyses of samples yielded respective precisions of 0.2 ‰ for 15NNO3 and 0.4

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‰ for 18ONO3.

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NO2- N and O isotope ratios (15NNO2 vs. air and 18ONO2 vs. V-SMOW) were measured at

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UConn by conversion to N2O with azide and analyzed on the GC-IRMS.35 Samples were measured

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in replicate batch analyses, standardized with NO2- reference materials N-23, N-7373 and N-

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1021920, as well as with internal standards (WILIS 10, 11, and 20; provided by Scott Wankel,

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WHOI, Woods Hole, MA) dissolved in high purity water at pH 10. For each batch analysis, some

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standards and representative samples were also diluted in

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isotopic exchange during conversion of NO2- to N2O by azide. Replicate analyses yielded

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respective precisions of 0.3 ‰ for 15NNO2 and 0.6 ‰ for 18ONO2. Following inter-calibration of

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NO2- isotope measurements with the NO2--specific denitrifier method36, we observed systematic

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offsets of 15NNO2 values in Experiment 2 (Fig. S1), which we attribute to a salinity effect of

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brackish samples not accounted for by our standards (Section S1; Fig. S1). We adjusted NO2- 15N

153

values accordingly. Values for 18ONO2 were not adjusted as isotopic offsets were negligible.

18O-enriched

water to account for

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The O isotope composition of water (18OH2O vs. V-SMOW) used in the incubations was

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measured at USGS using the CO2 equilibration method37, with typical precision of ±0.1‰

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(https://isotopes.usgs.gov/).

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Regression analyses

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Type II regressions38,39 of 18ONO2 vs. 18OH2O, 18ONO3 vs. 18OH2O, and closed-system Rayleigh

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plots for determination of isotope effects1 were computed in Matlab (Mathworks) with a script

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provided by Edward Pelter.40

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Figure 2. Results for Experiment 2: (A) Time-dependent evolution of [NH4+], [NO2-], and [NO3-] for all incubations following NH4Cl addition. (B) Time-dependent and (C) reaction progress-dependent evolution of 18O for NO2-. (D) Time-dependent and (E) reaction progress-dependent evolution of 18O for total NO3- (produced + background). (F) Time-dependent and (G) reaction progress-dependent evolution of 18O for NO3- produced. (H) Time-dependent and (I) reaction-progress-dependent evolution of 15N for NO2-. (J) Time-dependent and (K) reaction-progressdependent evolution of 15N for total NO3-. (L) Time-dependent and (M) reaction-progress- dependent evolution of 15N for NO3- produced. The respective terms fNOx are the fraction of NOx at the time of sampling relative to the maximum observed concentration of corresponding NOx. Dashed, colored lines correspond to hypothetical 18O of NO2- (panels B and C) and NO3- (panels F and G) for each 18OH2O treatment estimated from Eq. 1 and 2. Error bars are the analytical error in all panels with the exception of F, G, L, and M (18O and 15N NO3- produced), in which the errors were propagated from analytical uncertainty of total minus initial 18O and 15N NO3-.

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Results

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Observations from the preliminary fresh-water Experiment 1, which are analogous to those

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in brackish-water Experiment 2, are largely described in Supplementary Materials (Sections S2,

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Fig. S2-S4), in order to simplify the main text. Experiment 1 is less well resolved than Experiment

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2 because it had less 18OH2O variation and longer sampling intervals. Nevertheless, the patterns

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revealed therein provide valuable confirmation that the conclusions gleaned from Experiment 2

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are relevant to other aquatic systems. Results from Experiment 2 are described and discussed

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below.

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Time-dependent evolution of DIN.

171

The brackish stream water used for Experiment 2 had initial NO3- and NH4+ concentrations

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of 1.4 M and 4.8 M, respectively, with no detectable NO2- (< 0.1 µM). Following NH4Cl addition,

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initial experimental NH4+ concentrations were ~62 M (Fig. 2A). NH4+ increased by 7– 10 M in

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all bottles over the first ten days of incubation, with no NO2- or NO3- production, suggesting

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ammonification of dissolved organic nitrogen (DON). NO2- production was first detected on day

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14 and NO2- accumulated rapidly to 67 – 75 M, then more slowly to peak concentrations of 72

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– 78 M by day 33, at which time NO2- concentrations exceeded peak NH4+ concentrations by an

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average of 2.1 µM, suggesting that some ammonification of DON occurred concurrently with NH3

179

oxidation. NO3- production was first detected between day 38 and 43 among experimental

180

bottles, reaching maximum concentrations of 71 – 80 µM NO3- between days 43 and 64. The

181

amount of NO3- produced during the incubations ranged between 70 M and 79 M, 10 – 17 µM

182

in excess of initial NH4+ concentrations, consistent with concurrent ammonification of DON.

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183

Experiment 1 showed an analogous progression of DIN pools, albeit, with a more rapid onset of

184

NO2- production, followed by the immediate onset of NO3- production following the accumulation

185

of NO2- (Section S2.1; Fig. S2).

186

Evolution of 18O of NO2- and NO3-

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18ONO2, first measured on day 17, ranged between -6.2 ‰ and 18.9 ‰ in treatments 1 to 4,

188

respectively, when NO2- concentrations were 10 – 17 M, corresponding to an fNO2 value of 0.15

189

– where we define fNO2 as the measured NO2- concentration at the time of sampling relative to

190

the maximum NO2- concentration observed in each incubation, [NO2-]/[NO2-max], prior to the

191

onset of NO2- oxidation (Fig. 2B, 2C). 18ONO2 in all incubations then increased rapidly by ~8 ‰ as

192

NO2- concentrations increased concurrently to 67 – 75 M (day 23), increasing more slowly

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thereafter to 3.9 ‰ to 34.4 ‰ for treatments 1 – 4 at respective [NO2-] maxima (fNO2= 1; around

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day 40 in all but one incubation). These final 18ONO2 values produced from NH3 oxidation were

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notably different than those estimated from the weighted sum of 18O sources (Eq. 1; Fig. 2B,

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2C). Following the onset of NO3- production, 18ONO2 values continued to increase slightly, then

197

began to decrease after approximately half of the NO2- had been oxidized to NO3-.

198

18ONO3,total values increased rapidly at the onset of NO3- production, reflecting mixing of

199

newly produced NO3- with the modest background pool of stream water NO3- (1.4 M at 5.5 ‰;

200

Fig. 2D, 2E). 18ONO3,total values then evolved gradually as NO2- oxidation proceeded, posting final

201

values of 0.1 ‰ to 32.9 ‰ in treatments 1 to 4, respectively. The corresponding values of

202

18ONO3,produced gradually decreased from the onset of NO2- oxidation to final values (fNO3 = 1;

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where fNO3 = [NO3-produced]/[NO3-max produced]) of 0.0 ‰ to 33.6 ‰ for treatments 1 to 4, respectively.

204

The final 18ONO3,produced values differed from those estimated from Eq. 2 (Fig. 2F, 2G).

205

Experiment 1 values of 18ONO2 and of 18ONO3,produced showed parallel trends as in Experiment

206

2 (Section S2.2; Fig. S2).

207

Evolution of 15N of NO2- and NO3-

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The 15N of the NO2- produced initially (0.05 < fNO2 < 0.11; Fig. 2H, 2I) was -23.2 ± 0.4 ‰ among

209

treatments, and increased to a maximum of 7.3 ± 0.1 ‰ at peak NO2- concentrations on day 40

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(fNO2 = 1.0). As NO2- was subsequently depleted, 15NNO2 decreased progressively to -9.1 ‰ on

211

day 51 (fNO2 ~ 0.10), before NO2- was completely consumed.

212

The stream water NO3- had an initial 15NNO3 value of 10.8 ± 0.2 ‰ (Fig. 2J, 2K). 15NNO3,total

213

values rapidly increased at the onset of NO2- oxidation to an average value of 16.2 ± 0.9 ‰ on

214

day 48 (fNO3 ~ 0.10) in all treatments, before gradually decreasing to a final value of 7.8 ± 0.1 ‰

215

(Fig. 2J, 2K). Values of 15NNO3,produced at the onset of NO3- production (fNO3 < 0.10) were ~ 18 ‰

216

then decreased gradually as NO3- accumulated to a final value (fNO3 = 1) of 7.7 ± 0.1 ‰ among

217

treatments (Fig. 2L, 2M). Similar dynamics for 15NNO2 and 15NNO3 were observed in Experiment

218

1 (Section S2.3; Fig. S3)

219

Dependence of 18ONO2 and 18ONO3 on 18OH2O

220

The values of 18ONO2 at peak NO2- prior to the onset of NO3- production correlated linearly

221

to 18OH2O among treatments (Fig 3A). The slope of the linear regression was 0.73 ± 0.0, greater

222

than the value of 0.50 expected from Eq. 1, thus signaling a greater dependence of 18ONO2 on

223

18OH2O. The intercept, 5.2 ± 0.0 ‰, was lower than the value of 12 ‰ (one half of 18OO2)

13



224

expected from Eq. 1, assuming a 18OO2 of 24 ‰ for O2 in equilibrium with the atmosphere.

225

Similarly, final values of 18ONO3,produced also show a linear dependence on 18OH2O, with a slope

226

of 0.93 ± 0.01 and intercept of 2.0 ± 0.1 (Fig 3B); that is, a regression slope higher than the value

227

of 0.66 and an intercept lower than the value of 8 ‰ that would result from unfractionated O

228

contributions from O2 and H2O (Eq. 2). Similar dependencies of 18ONO2 and of 18ONO3,produced on

229

18OH2O were observed in Experiment 1 (Section S2.4; Fig. S4).

230 231 232 233

Figure 3. (A) 18ONO2 vs. 18OH2O at peak [NO2-] prior to NO3- production and (B) final 18ONO3, 18 produced vs.  OH2O. Equations on bottom correspond to the linear fit of the observations. Markers are larger than error bars, based on the standard deviation of triplicate bottle incubations.

18O dynamics during NH3 oxidation to NO2-

234

The 18O of NO2- produced by nitrification in our incubations, in both Experiments 1 and 2,

235

does not conform to expectations from Eq. 1. If the O atoms incorporated during the oxidation

236

of NH3 to NO2- were derived equally from molecular O2 and H2O,8 without isotopic fractionation,

237

and without NO2-H2O O atom exchange, then the 18O of NO2- produced by nitrification in our

238

incubations would correspond to the weighted sum of the end member 18O values of the

239

reactants (Eq. 1; Fig. 3A). Further, the 18O of NO2- produced would remain invariant throughout

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the time course of the experiment, provided the 18O of the O2 and H2O substrate pools remained

241

constant. Contrary to these expectations, the 18O of NO2- produced at the onset of NH3 oxidation

242

was substantially lower (by 10 ‰ to 17 ‰ among treatments) than predicted by the weighted

243

sum of the 18O source values (Eq. 1; Fig. 2B, 2C; Fig. S2), even lower than the 18O of water. The

244

low 18O of NO2- values initially produced indicate isotopic discrimination during O-atom

245

incorporation from either O2, H2O, or both. As NH3 oxidation proceeded, 18ONO2 values increased

246

progressively, reaching 18ONO2 maxima at peak NO2- concentrations. This pattern is consistent

247

with NO2- isotopic exchange with H2O, whereby the 18ONO2 values increased to values

248

approaching isotopic equilibrium (18ONO2 = 18OH2O + 13 ‰20,41). Without exchange, 18ONO2

249

values- would have remained relatively low and constant due to isotopic discrimination with

250

respect to the 18O values of the larger O2 and H2O substrate pools. Nevertheless, at peak NO2-

251

concentrations in all incubations, 18ONO2 values remained lower than expected for full

252

equilibration in each treatment (with different 18OH2O values), indicating O exchange between

253

NO2- and H2O was incomplete among treatments.

254

These observations support a more complex model of NH3 oxidation that includes isotopic

255

effects of oxygen atom incorporation from molecular O2 and H2O, as well as partial oxygen

256

isotopic equilibration of NO2- with H2O (Fig. 118,19):

257

δ18ONO2 = [2 ( 18OO2 ― 18εk,O2)+ 2 (18OH2O ―

258

+ (18OH2O + 18εeq)(XNO2,1)

1

1

18ε k,H2O,1)]

(1 – XNO2,1)

(3a)

259

where XNO2,1 is the fraction of NO2- oxygen atoms that have exchanged with H2O during NO2-

260

production, 18k,O2 is the kinetic isotope effect for O2 incorporation, 18k,H2O,1 is the kinetic isotope

261

effect for H2O incorporation, and 18eq is the equilibrium isotope fractionation factor between 15



Page 16 of 35

262

NO2- and H2O. Equation 3a can be re-arranged to a linear model describing 18ONO2 vs. 18OH2O18,19

263

(Fig. 3A):

264

δ18ONO2 = (2 + 2 XNO2,1)18OH2O

265

1

1

(3b)

1

+ 2[(δ18OO2 ― 18εk,O2 ― 18εk,H2O,1)(1-XNO2,1)] + ( XNO2,118εeq)

266

From Eq. 3b compared to the slope of the linear regression of peak NO2- 18ONO2 vs. 18OH2O (Fig.

267

3A), we estimate the fraction of NO2- O atoms that exchanged with H2O prior to the onset of NO3-

268

production, XNO2,1, to be 47 ± 1 % (Table 2). The combined O kinetic isotope effects during NH3

269

oxidation, 18k,O2 + 18k, H2O,1, are computed from the regression intercept, assuming a 18OO2 value

270

of 24 ‰7 and a 18eq value of 13 ‰ for both abiotic and biologically mediated equilibration20,41,

271

resulting in 18k, O2 + 18k, H2O,1 = 27.3 ± 4.0 ‰ (Table 2). The derived XNO2,1 in Experiment 1 was

272

lower, 26 ± 27 %, though less certain, as was 18k,O2 +18k,H2O,1, at 7.5 ± 17.8 ‰ (Table 2; Section

273

S2.4). We note that simple mass-balance calculations indicate that potential increases in 18OO2

274

from coincident respiration in the incubations are negligible, and thus do not explain observed

275

increases in 18ONO2 (Section S3).

Table 2. Oxygen isotope effects and fraction of NO2- isotopically equilibrated with H2O (XNO2) in the incubations, compared to previously reported estimates from nitrifier culture studies and sea water incubations.18,19,21 ‡Reported values correspond strictly to biologically mediated equilibration. *The computed kinetic isotope effect may be less negative than the true effect because it includes the competing effect of partial NO2-H2O equilibration. ND = no data.

16


Page 17 of 35


Seawater CoIncubations cultures

NH3 oxidizing cultures

NO2oxidizing cultures

Experiment 1

Experiment 2

(n = 1)

(n = 1)

(n = 7)19

(n = 6)19

(n = 10)18

(n = 3)21

XNO2,1 (%)

26 ± 27

47 ± 1

35 – 100

16 - 28‡

1 – 25‡

-

XNO2,T (%)

74 ± 12

78 ± 3

48 – 100

0 - 26‡

-

0 – 3‡

7.5 ± 17.8

27.3± 4.0

11 – 20

16 – 23

18 – 30

-

22.5 ± 6.5

13.5 ± 2.0

1 – 27

6 – 12

-

9 – 25

ND

ND

-

-10 – -1.4

18

k,O2

(‰) 18

+18k,H2O,1

k,H2O,2

18

k,NO2

(‰)

(‰)

ND

[-3.9 ±

0.3]*

276 277

18O dynamics during NO2- oxidation to NO3-

278

During biological NO2- oxidation, the O atom appended to NO2- is reported to derive from

279

H2O.9 In the absence of isotopic fractionation, the 18O of NO3- produced during nitrification

280

would then reflect the weighted sum of the end member 18O values of the reactants (Eq. 2), and

281

would remain relatively constant within individual treatments throughout the experiment.

282

Instead, observed values of 18ONO3,produced decreased continually among all treatments following

283

onset of NO3- production (Fig. 2F, 2G). This trend is consistent with inverse kinetic isotopic

284

discrimination on the O atoms of NO2- during its oxidation (18k,NO2), which has been documented

285

in culture studies of NO2- oxidizing bacteria.21,42 Inverse isotopic fractionation of O is also

286

apparent from the corresponding 18ONO2 values, which decreased progressively as NO2- was

287

oxidized to NO3-. The inverse isotope effect can be approximated from the closed-system

288

Rayleigh substrate equation1,

289

𝛿𝑠𝑢𝑏𝑠𝑡𝑟𝑎𝑡𝑒 = 𝛿𝑠𝑢𝑏𝑠𝑡𝑟𝑎𝑡𝑒,

𝑖𝑛𝑖𝑡𝑖𝑎𝑙

17

― ε × ln (𝑓)


(4)


Page 18 of 35

290

yielding a value of -3.9 ‰ for 18k,NO2 (Fig. S5; Table 2). We note, however, that this value is likely

291

less negative than the actual fractionation factor given a competing influence of ongoing NO2-

292

equilibration that tended to increase 18ONO2 in all treatments, and given the concurrent

293

ammonification of DON that likely contributed to an additional progressive increase in NO2-

294

concentrations. In Experiment 1, inverse isotopic discrimination of NO2- is also apparent (Fig. S2),

295

but the poor time resolution precludes derivation of a value for 18k,NO2.

296

In order to determine whether NO2- was subject to further O atom exchange following the

297

onset of NO3- production, we investigated the correspondence of NO3- O isotope dynamics to the

298

more comprehensive description of NO2- oxidation that accounts for equilibration as well as O

299

incorporation kinetic isotope effects19,21.

300

δ18ONO3, produced = 3 [(1 ― XNO2,T)18ONO2 + XNO2,T(18OH2O+ 18εeq)] + 3 (18OH2O ―

301

The exchange term, XNO2,T, accounts for the total fraction of NO2- equilibration for the duration

302

of our incubations, and 18εk,H2O,2 is the kinetic isotope effect associated with O atom incorporation

303

from water into NO2- during oxidation to NO3-. As presented, Eq. 5a neglects the inverse kinetic

304

isotope effect for NO2- conversion to NO3-, as this effect would be minimized by the end of the

305

experiments after complete oxidation of NO2- to NO3-. It would, however, be pertinent in the case

306

of incomplete NO2- oxidation. With the substitution of the 18ONO2 term from Eq. 3, Eq. 5a is

307

rearranged to conform to a linear formulation of 18ONO3, produced vs. 18OH2O19,21.

308

δ18ONO3, produced = (3 + 3 XNO2,T)18OH2O

309 310

2

1

2

18ε k,H2O,2)

1

(5a)

(5b)

1

+ 3[(δ18OO2 ― 18εk,O2 ― 18εk,H2O,1)(1-XNO2,T) ―

18ε k,H2O,2)]

2

+ 3 ( XNO2,T18εeq)

The resulting XNO2,T , derived from the regression slope (Fig. 3B), is ~ 78%. 18


Page 19 of 35


311

Eq. 5b also permits derivation of the kinetic isotope effect for O incorporation from water

312

during NO2- oxidation, 18εk,H2O,2, computed from the intercept of 18ONO3, produced vs. 18OH2O (Fig.

313

3B), yielding a value of 13.5 ± 2.0 ‰ (Table 2).

314

In Experiment 1, derived values for XNO2,T and 18εk,H2O,2 were similar to those in Experiment 2,

315

74 ± 12% and 22.5 ± 6.5‰, respectively (Table 2; Section S2.4).

316

Kinetic N isotope fractionation during NH3 and NO2- oxidation

317

Although N isotope effects were not the focus of this study, the temporal resolution of

318

Experiment 2 allows for the derivation of apparent kinetic N isotope fractionation factors for

319

nitrification. The gradual increase of 15NNO2 during NO2- production (Fig. 2H, 2I) is consistent with

320

a kinetic isotope effect on N during NH3 oxidation. 15NNO2 values were modeled with a closed-

321

system Rayleigh product equation1 (Eq. 6; Fig. 4A):

322

δproduct = δsubstrate,

initial

―ε ×

𝑓ln (𝑓) (1 ― 𝑓)

(6)

323

where f is fraction of the substrate remaining relative to the initial substrate concentration;

324

results indicate a N isotope fractionation factor for NH3 oxidation (15k,AMO) of 38.9 ± 0.5 ‰. We

325

note that this value may be slightly depressed due to coincident ammonification. We also note

326

this approach includes the assumption that NH4+ was not concurrently assimilated for anabolism

327

by heterotrophic microbes, given the evident catabolic ammonification of DON, such that

328

isotopic discrimination of NH4+ derives from its nitrification only.

329

In addition, both substrate (Eq. 4) and product (Eq. 6) equations for a closed-system Rayleigh

330

model1 yield reassuringly similar estimates of the kinetic N isotope effect during NO2- oxidation

331

of 15k,NO2 of -9.5 ± 0.2‰ and -10.3 ± 0.4 ‰ (Fig. 4B, 4C).

19



332

Figure 4. Closed-system Rayleigh plots1 of 15N in NO2- and NO3- in Experiment 2, based on Eq. 4 and 6: (A) 15NNO2 vs. f*ln(f)/(1-f) during NH3 oxidation, (B) 15NNO2 vs. ln (f) during NO2- oxidation, and (C) 15NNO3 vs. f*ln(f)/(1-f) during NO2- oxidation. All 12 incubation bottles are represented. Kinetic isotope effects correspond to respective slopes of fitted linear regressions. Error bars correspond to the standard deviation of replicate measurements.

333

Discussion

334

Oxygen isotope effects and NO2- isotopic equilibration during nitrification

335

The oxygen isotopic compositions of NO2- and NO3- produced from the biological oxidation of

336

NH3 in our stream water experiments were directly related to 18OH2O among treatments, but the

337

relations were different from those implied by weighted 18O contributions of unfractionated O

338

atom sources (Eq. 1 and 2). The divergence from this simple source attribution suggests that

339

kinetic isotope effects on O during its incorporation were occurring, as well as O atom

340

equilibration of NO2- with H2O with an associated equilibrium isotope effect, as documented for

341

cultures of nitrifying prokaryotes.18,19,21 Our observations can be explained with an isotope mass

342

balance model that accounts for these dynamics (Eq. 3 and 5). The kinetic O isotope effects thus

343

derived are similar to those observed in cultures of nitrifying prokaryotes and in seawater

344

incubations (Table 2), suggesting that the O isotopic systematics of nitrification are homologous

345

among ecosystems.

20


Page 20 of 35

Page 21 of 35


346

The proportions of NO2- O atoms exchanged with H2O throughout the incubations are similar

347

to those observed during seawater incubations, but greater than those observed in NH3 and NO2-

348

oxidizing cultures (Table 2).18,19 We note, however, that our estimates of XNO2,1 and XNO2,T

349

represent combined effects of abiotic and biologically-mediated equilibration, in contrast to

350

estimates from cultures that correspond to the fraction of biologically-mediated exchange

351

exclusively (Table 2). Isotopic exchange of O atoms between NO2- and H2O occurs abiotically, as

352

a function of pH and temperature,20,41 and can also be facilitated during nitrification by NH3

353

oxidizing bacteria and archaea.18,19 At the pH of ~8.1 in our incubations and room temperature,

354

the expected abiotic equilibration rate constant (kabiotic) is approximately 0.02 per day.41 We

355

obtained similar values from Experiment 2 during the ~3-week period between the end of NO2-

356

production and the onset of NO2- consumption, averaging 0.018 ± 0.003 day-1 among treatments.

357

Preceding this interval, the isotopic equilibration in our incubations (in Experiment 2) was

358

substantially more rapid, with k averaging 0.090 ± 0.024 day-1. The steep 18ONO2 increase that

359

occurred concurrently with NO2- production is interpreted as predominantly biologically-

360

mediated exchange with rate constant (kbio) of approximately 0.073 day-1, 4-fold greater than

361

kabiotic. The rate of biologically mediated equilibration is evidently linked to the rate of NH3

362

oxidation and may be facilitated by the acidic pH of the cellular periplasm, wherein the

363

conversion of the hydroxylamine intermediate to NO2- is catalyzed.

364

In contrast, biologically-mediated O atom exchange by NO2- oxidizing bacteria is thought to

365

be negligible (Table 2)21, such that the increase in the fraction of NO2- exchanged in Experiment

366

2, from 47% at the onset of NO3- production to 78% at the end of the incubations, likely derives

367

dominantly from abiotic equilibration. Indeed, the rate of exchange during NO2- oxidation in our

21



368

experiment (k = 0.02 day-1) is consistent with the estimated abiotic equilibration rate constant

369

expected for our pH and temperature conditions (kabiotic = 0.02 day-1). In summary, the total NO2-

370

isotopic exchange in the incubations integrated both biologically-mediated exchange occurring

371

specifically during ammonium oxidation and abiotic exchange occurring throughout the whole of

372

the incubations. The observed exchange is comparable to that reported for incubations of natural

373

seawater assemblages (Table 2), where high concentrations of accumulated NO2- (upwards of 50

374

M) remained present for at least 13 days prior to NO2- oxidation.19 Exchange was greater,

375

however, than that observed in mono-cultures or in co-cultures, in which NO2- did not

376

accumulate and for which biotic equilibration is reported specifically (Table 2).

377

N isotope effects for NH3 and NO2- oxidation were also evident in the incubations. The

378

magnitudes of these effects were larger than those reported previously for cultures and marine

379

isolates, which generally range between 12 ‰ and 20 ‰, but they were similar to those reported

380

for freshwater bacterial isolates such as Nitrosomonas europaea and Nitrosomonas eutropha.1,43

381

To the extent that larger N isotope effects may be diagnostic of freshwater nitrifiers, NH3 (and

382

NO2-) oxidation in the brackish water of Experiment 2 may have been conducted largely by

383

freshwater nitrifiers, over potentially less abundant or less active marine nitrifiers, conceivably

384

explaining the considerable lag phases preceding the respective onsets of NH3 and NO2- oxidation

385

compared to Experiment 1. The amplitude of the inverse N isotope effect during NO2- oxidation

386

(in Experiment 2) was, however, comparable to that observed in cultures of a marine NO2-

387

oxidizing bacterium, between -21 and -9‰21.

388

Implications for interpretation of NO3- isotope ratios in terrestrial environments

22


Page 22 of 35

Page 23 of 35


389

The observations in this study are congruent with some observations that 18O values

390

produced by nitrification in soils are often lower than would be expected from relative source

391

contributions, such that O equilibration and/or substantial O kinetic isotope effects must be

392

invoked to account for these observations (Fig. 5A).14,15,44 Our experiments are also consistent

393

with an empirical convention adopted by the oceanographic community that the 18O of nitrified

394

NO3- in marine systems is on the order of 18OH2O + 1 ‰ (Fig 5A).19,29,45–49 Indeed, a density

395

distribution of 3,360 plausible nitrification 18ONO3 values predicted by Eq. 5 illustrates that over

396

60 % of solutions are within ±3 ‰ of ambient 18OH2O (Fig. 5B). We computed these values for a

397

18O H2O of -5‰ using reported ranges of kinetic isotope effects for O in increments of 1 ‰

398

(Table 2) and XNO2 values of 0 %, 25 %, 50 %, 75 %, and 100 %. We thus propose that empirically,

399

the 18ONO3 values produced by nitrification of NH3 in saturated experimental systems and marine

Figure 5. (A) 18O of nitrification NO3- vs. corresponding 18O of H2O observed here and in other studies using comparable methods and calibrations. Shapes of symbols indicate categories of samples, described in Table S1. Respective dotted lines correspond to expected 18ONO3 values from Eq. 2 and to the corresponding 18OH2O. (B) Density distribution of NO3- 18O values expected for nitrification at a 18O H2O of -5‰ based on observed ranges of kinetic isotope effects (Table 2) and XNO2 values of 0%, 25%, 50%, 75%, and 100%. The dashed line represents 18O of nitrification NO3- predicted by Eq. 2.

23



400

ecosystems are more akin to the 18O of water than to values expected from weighted 18O of

401

O-atom sources (Fig. 5B).

402

Although centered on the 18O of water, potential nitrification 18ONO3 values from this and

403

other studies cover a relatively broad range (Fig. 5B). Among influential parameters, 18ONO3

404

values are particularly sensitive to the extent of NO2- isotopic exchange, a poorly constrained

405

term in the environment.15,50 Biologically-mediated exchange of NO2- with water during NH3

406

oxidation appears to converge on a proportion of 0-25 % (Table 2). The rate of abiotic exchange,

407

in turn, increases with temperature and acidity, such that complete equilibration takes weeks to

408

months at near neutral pH, and is on the order of hours to days at below-neutral pH.51 In this

409

respect, the pH of some soils and groundwaters can be sufficiently low (between 4 and 7) to

410

promote relatively rapid abiotic NO2--H2O equilibration, including some well-drained agricultural

411

areas with high NO3- concentrations.52,53 Fang et al. (2012)14 observed increasing soil 18ONO3

412

values with elevation, which they attribute to a higher proportion of NO2- isotopic exchange with

413

increasing soil acidity. The extent to which NO2- exchanges abiotically with H2O is a function of its

414

residence time, which can range from minutes in some systems to days and weeks or more in

415

others.20,36,41,54,55 The influence of the NO2- residence time on its isotopic exchange is evident in

416

our study, as well as other nitrification incubations in which NH3 oxidation and NO2- oxidation

417

were partially decoupled, allowing for high concentrations of NO2- to accumulate and an increase

418

in the amount of abiotic exchange that occurs.15,19 Among incubations of different soil types,

419

Snider et al. (2010)15 observed between 37 % and 52 % isotopic exchange with water, which was

420

inversely correlated with net NO3- production rates, as soils with slower nitrification rates showed

24


Page 24 of 35

Page 25 of 35


421

a higher propensity for NO2- accumulation. In all, the NO2- isotopic exchange term associated with

422

nitrification is likely to differ among soil types, as a function of pH and organic material content.15

423

Despite considerable variability in the spectrum of model results based on experiments, none

424

of the solutions in the density distribution (Fig. 5B) are as high as the 18ONO3 value predicted by

425

Eq. 2 (i.e., +4.7 ‰). Yet, numerous observations in soils and oxic groundwaters appear to

426

converge on values closer to those predicted by Eq. 2, as illustrated in a partial compilation of

427

comparative data representing varying land use, climate, and NO3- concentrations (Fig. 5A; Table

428

S1), omitting anomalously high reported 18ONO3 values that may have been affected by known

429

analytical artifacts16 (Section S4), and including only groundwater samples that lack dissolved-gas

430

evidence of denitrification within the saturated zone56,57 or partially saturated infiltration

431

water.58 Some potential reasons why these 18ONO3 values are elevated above experimental

432

nitrification values include: (1) Components of NO3- with high 18O from atmospheric deposition

433

(60-90 ‰) or artificial NO3- fertilizer (20-30 ‰) may be mixed with nitrification-derived NO3- in

434

soils and infiltrating pore waters. Un-cycled atmospheric or fertilizer-sourced NO3- may be

435

important locally (e.g., in arid regions or agricultural areas), but probably not uniformly across

436

the range of land uses, climate conditions, and nitrate concentrations represented by

437

groundwater samples (Fig. 5A). (2) Because of local or transient evaporation or isotopic variation

438

of precipitation, 18OH2O values of groundwater or bulk soil water may not be representative of

439

values at nitrification sites, which could be higher; however, these effects would not be expected

440

to yield consistent offsets across climate and land use gradients. (3) Similarly, local or transient

441

O2 respiration causes isotopic fractionation, such that extractable bulk 18OO2 values may not be

442

representative of values at nitrification sites. However, the influence of the isotopic composition

25



443

of O2 on NO3--nit is relatively small and decreases as a function of NO2- isotopic exchange. For

444

instance, even with no exchange of NO2-, ambient 18OO2 values would need to exceed 60 ‰,

445

corresponding to a closed-system O2 pool that is ~85 percent respired1,59, in order for NO3--nit to

446

fortuitously conform to the 18ONO3 expected from Eq. 2, given a 18OH2O of -5 ‰ and median O

447

isotope fractionation effects. (4) Soil NO2- may consistently fully equilibrate with H2O before

448

being oxidized to NO3-, thus resulting in relatively high 18ONO3, as indicated by some of the

449

highest values in Fig. 4B (Results of models with XNO2 = 100% range from -5.7 ‰ and +2.0 ‰);

450

however, this would be inconsistent with some direct observations of the extent of NO2- isotopic

451

exchange15 and it would not account for a preponderance of values near that indicated by Eq. 2.

452

To explain apparent conformity to simple source attribution of O atoms more

453

comprehensively across terrestrial systems, we consider that (5) denitrification in anaerobic

454

microsites can increase the 18ONO3 values otherwise ascribed to nitrification, in which case a

455

comparable increase in 15NNO3 should also be evident.46,56,57,60,61 (6) Finally, rapid redox cycling

456

between NO3- and NO2- can increase 18ONO3 relative to 15NNO3 values51,62, by not allowing for

457

isotopic equilibration of elevated 18ONO2 values imparted by the branching isotope effect during

458

NO3- reduction (aka, intra-molecular isotope effect).20,29,63,64 The NO3- produced by re-oxidation

459

of high 18O NO2- then has a greater 18ONO3 than that originally reduced, but a similar 15NNO3.50

460

More measurements of NO2- 18O values in various environments could be useful for evaluating

461

this potential dynamic.

462

In contrast to soils, water-saturated terrestrial environments where nitrification of NH3 is a

463

relatively discrete uni-directional process may be more likely to produce nitrification 18ONO3

464

values close to those of ambient water, as in oxygenated marine systems. For instance, at the

26


Page 26 of 35

Page 27 of 35


465

down-gradient end of a groundwater flow-through lake in Massachusetts, nitrification occurred

466

in the subsurface as oxic lake water infiltrated into the underlying aquifer.65 NO3- produced during

467

infiltration had 18ONO3 = -5.4 ± 0.5 ‰ in an interval where 18OH2O = -5.6 ± 0.1 ‰ and 18OO2 =

468

22-26 ‰ (Fig. 5A- X symbols; Table S1). These lakebed porewater data are consistent with the

469

saturated experiments and marine data, but distinctly different from most of the soil NO3- and

470

soil-derived oxic groundwater NO3- data, possibly indicating fundamental differences between

471

these nitrification environments. Some evidence exists that nitrification in partially saturated

472

wastewater infiltration systems might generate intermediate 18ONO3 values.58

473

In summary, 18ONO3 values produced from the nitrification of NH3 do not conform to a simple

474

O source attribution model (Eq. 2). Rather, NO3- originating from the direct nitrification of NH3

475

acquires a 18O value close to that of H2O. Nevertheless, substantial variability in the 18ONO3

476

associated with nitrified NO3- is expected. Notwithstanding inherent differences in kinetic isotope

477

effects among nitrifier communities, the degree of NO2- isotopic exchange is likely dependent on

478

soil/water properties, compelling evaluation of the propensity for NO2- isotopic equilibration

479

among environments. While numerous field observations present a fortuitous match of nitrified

480

NO3- 18O to the simple source attribution model, we propose that co-incident denitrification or

481

repeated redox cycling between NO3- and NO2- pools could underlie such trends. Coupled N and

482

O isotope ratio measurements of NO3- and NO2- pools can provide diagnostic tracers to elucidate

483

these dynamics in situ, to arrive at a more robust understanding of biological N cycling and,

484

ultimately, more accurate diagnosis of natural and pollutant sources of reactive nitrogen to the

485

environment.

486

27



487

Supporting Information. Description of 15N and 18O NO2- corrections, preliminary

488

Experiment 1 results, determination of the lack of influence of respiration on 18O of dissolved

489

O2, discussion of artifacts associated with different 18ONO3 analytical methods.

490

Acknowledgments. This study was supported by a grant from National Science Foundation

491

Division of Earth Sciences (EAR-1252089) to J.G. and C.T., and by the U.S. Geological Survey

492

(USGS) Water and Environmental Health Mission Areas. Assistance in the laboratory was

493

provided by Claudia Koerting and Reide Jacksin at the University of Connecticut and Janet Hannon

494

and Stanley Mroczkowski at USGS. We would like to thank USGS reviewer Christopher Green

495

and three anonymous reviewers for their constructive comments. Any use of trade, firm, or

496

product names is for descriptive purposes only and does not imply endorsement by the U.S.

497

Government.

28


Page 28 of 35

Page 29 of 35


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(2)

Ward, B. The global nitrogen cycle. In Fundamentals of geobiology; Blackwell Publishing Ltd.: Hoboken, 2012; pp 36–48.

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Daims, H.; Lebedeva, E. V.; Pjevac, P.; Han, P.; Herbold, C.; Albertsen, M.; Jehmlich, N.; Palatinszky, M.; Vierheilig, J.; Bulaev, A.; et al. Complete nitrification by Nitrospira bacteria. Nature 2015, 528 (7583), 504–509.

(4)

Kendall, C.; Elliott, E. M.; Wankel, S. D. Tracing anthropogenic inputs of nitrogen to ecosystems, Chapter 12. In Stable Isotopes in Ecology and Environmental Science; Blackwell Publishing Ltd.: Hoboken, 2007; pp 375–449.

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Barkan, E.; Luz, B. High precision measurements of 17O/16O and 18O/16O ratios in H2O. Rapid Commun. Mass Spectrom. 2005, 19 (24), 3737–3742.

(6)

Benson, B. B.; Krause, Daniel, J. The concentration and isotopic fractionation of gases dissolved in freshwater in equilibrium with the atmosphere. 1. Oxygen. Limnol. Oceanogr. 1980, 25 (4), 662–671.

(7)

Kroopnick, K.; Craig, H. Atmospheric oxygen: isotopic composition and solubility fractionation. Science (80-. ). 1972, 175 (4017), 54–55.

(8)

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