# δ2H of water from fluid inclusions in Proterozoic halite: Evidence for a deuterium-depleted hydrosphere?

Christophe Lécuyer, François Fourel, Nigel Blamey, Uwe Brand, Philip Fralick

► To cite this version:

Christophe Lécuyer, François Fourel, Nigel Blamey, Uwe Brand, Philip Fralick. δ2H of water from fluid inclusions in Proterozoic halite: Evidence for a deuterium-depleted hydrosphere?. Chemical Geology, 2020, 541, pp.119583. 10.1016/j.chemgeo.2020.119583. hal-02523013

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$\delta^{2}\text{H}$ of water from fluid inclusions in Proterozoic halite:

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Evidence for a deuterium-depleted hydrosphere?

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Christophe Lécuyer¹, François Fourel², Nigel Blamey³, Uwe Brand⁴, Philip Fralick⁵

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¹Laboratoire de Géologie de Lyon, UMR CNRS 5276, University of Lyon and Institut

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Universitaire de France, 69622, France. clecuyer@univ-lyon1.fr

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²Laboratoire d'Ecologie des Hydrosystèmes Naturels et Anthropisés LEHNA, UMR CNRS

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5023, University of Lyon, France. francois.fourel@univ-lyon1.fr

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³Department of Earth Sciences, University of Western Ontario, London, N6A 5B7, ON,

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Canada. nblamey2@uwo.ca

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⁴Department of Earth Sciences, Brock University, 1812 Sir Isaac Brock Way, St Catharines,

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ON L2S 3A1, Canada. ubrand@brocku.ca

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⁵Department of Geology, Lakehead University, Thunder Bay, ON P7B 5E1, Canada.

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pfralick@lakeheadu.ca

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Keywords: Proterozoic; water cycle; $^{2}\text{H}/^{1}\text{H}$ ratio; halite; fluid inclusion

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**Abstract** – $\delta^{2}$H were measured in water present in tiny primary fluid inclusions trapped in Mesoproterozoic (≈ 1.4 Ga) and Neoproterozoic (≈ 0.8 Ga) halite crystals. While the hydrogen concentrations range from 300 to 1500 ppm for the whole sample collection, $\delta^{2}$H values range from -74 ‰ to -54 ‰ (VSMOW) for the Mesoproterozoic halite sampled from the Sibley Group, Ontario, Canada, and from -89 to -38 ‰ for the Neoproterozoic halite from the Browne Formation, Officer Basin, Australia. The amount of evaporation required to precipitate halite is accompanied by a deuterium-enrichment of 30 ‰ to 90 ‰. It means that the original aqueous solutions had $\delta^{2}$H values comprised between -140 ‰ and -90 ‰.

Considering the estimated tropical paleolatitudes for both depositional environments, surface waters were most likely significantly deuterium-depleted relative to modern ones. If such interpretation is partly at variance with the geological record, it could be reconciliated with some modelling outputs of the long-term water cycle. Indeed, several mechanisms produce a progressive deuterium-enrichment of the oceans. Those identified so far are the addition of deuterium-rich extraterrestrial water (cometary dust), the loss of H during water photodissociation by ultraviolet radiation in the upper atmosphere, and the isotopic fractionation that takes place during the hydroxylation of nominally-anhydrous mantle rock-forming minerals. We also propose that the major process able to shift through time the $\delta^{2}$H of the hydrosphere is the reduction of H$_{2}$O into H$_{2}$ during the serpentinization of mantle rocks at mid-ocean ridges.

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# 1. Introduction

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Knowledge of the global water cycle is of paramount importance regarding the dynamics of mass transfer between the Earth's layers and the distribution of water among the main reservoirs. Indeed, the two main water reservoirs are the mantle and the oceans, while smaller but non-negligible amounts of freshwater are linked to the presence of continental masses (ice caps, glaciers, groundwaters, lakes). For decades, stable isotope compositions ($\delta^2$H and $\delta^{18}$O) of the water molecule have been recognized as useful tracers of the Earth's surface water cycle (Craig, 1961; Craig and Gordon, 1965) in relation to the heat balance that triggers climate change. Global scale processes related to Earth's internal dynamics have the potential to affect the isotopic composition of seawater. For instance, changes in seafloor spreading and subduction rates modify the rate of water transfer between the mantle and oceanic reservoirs. Changes in the depth of hydrothermal interaction between seawater and the oceanic lithosphere modify the amount of water that can be cycled into the mantle. The growth of geochemical reservoirs, like the continental crust or the ice caps, may also change the mass and isotopic compositions of the oceans.

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The knowledge of the stable isotope composition of the past hydrosphere represents a major issue as it reflects changes in the dynamics of the global water cycle. Therefore, the oxygen isotope composition of past oceans has been investigated through the study of ophiolites (Gregory and Taylor, 1981; Holmden and Muehlenbachs, 1993) and marine sediments (Veizer et al., 1997; Kasting et al., 2006), and was also the object of numerical modelling (e.g., Lécuyer and Allemand, 1999; Wallmann, 2001; Hren et al., 2009; Pope et al., 2012). The debate has been vigorous between the proponents of $^{18}$O-depleted (relative to SMOW) and “cold” Precambrian oceans and those supporting the theory of modern-like seawater $\delta^{18}$O and “warm” oceans. However, Came et al. (2007) argued for a constant $\delta^{18}$O of oceans throughout the Phanerozoic and, recently, Tartèse et al. (2017, page 55) reported
oxygen isotope compositions of carbonaceous matter indigenous to Precambrian cherts and
showed that “the oxygen isotope composition of seawater during most of the Precambrian
remained around 0±5‰ consistent with the composition of present-day seawater”. The main
process that maintained the steady-state δ¹⁸O value of the oceans is the buffering through
isotopic exchange between seawater and oceanic crust during hydrothermal alteration
(Muehlenbachs and Clayton, 1976; Gregory and Taylor, 1981).

By contrast, the δ²H of the oceans was likely not steady over time, especially during
the Precambrian, as supported by some geological data (Pope, 2011 and references therein)
and flux modelling outputs (Lécuyer et al., 1998; Genda and Ikoma, 2008; Kurokawa et al.,
2018). However, δ²H values of hydrothermally-altered past oceanic rocks (ophiolites,
greenstone belts) are tricky to interpret as they result from a combination of several variables
such as the temperature of isotopic exchange, the amount of water reacting with rock, and the
initial δ²H of reacting rock and water. Not least is the issue of the preservation of the
hydrogen isotope compositions of these old rocks. In most cases, hydrogen isotope
measurements were performed on serpentine minerals for which Kyser and Kerrich (1991)
showed that hydrous minerals are prone to retrograde isotopic exchange with water at low
temperature. This property may partly explain why serpentines of any geological age display
large ranges (≥ 50‰ in most cases) of δ²H values without showing any kind of secular trend
(see Figure 4.1 in Pope, 2011).

In the case of modelling, the magnitude and rates of isotopic changes are still poorly
constrained because of the large uncertainties associated with all the input and output fluxes
connecting the oceanic reservoir (Lécuyer et al., 1998; Genda and Ikoma, 2008; Kurokawa et
al., 2018). However, there is an emerging consensus toward the paradigm of a deuterium-
depleted Precambrian oceanic reservoir. Indeed, to a first approximation, we can consider that
the two main fluxes of water are regulated by plate tectonics. Seawater is incorporated in the
oceanic crust during hydrothermal alteration while it is partly released during dehydration of
subducted plates and hotspot volcanism (Kasting and Holm, 1992). Hydrogen isotopes
fractionate between free water and hydroxyl groups of minerals at all metamorphic
temperatures (Suzuoki and Epstein, 1976). The light isotope is preferentially incorporated
into the solid phase leaving the residual water enriched in deuterium. Mass balance and
kinetic treatments of the long-term water cycle suggested that the δ²H value of Archean
oceans could have been a few tens of permil lower than the present-day value (Lécuyer et al.,
1998; Kurokawa et al., 2018). This result could be explained by the imbalance between fluxes
of water trapped at ridges and released along subduction zones (Korenaga et al., 2017), by the
reduced size of continental masses (Condie and Kröner, 2013), and the absence of ice sheets
(Ramstein, 2011). Moreover, a progressive deuterium-enrichment of the oceans has been
proposed through an addition of water via impacts of extra-terrestrial bodies and the escape of
atomic H to space by photodissociation of the water in the upper atmosphere (Lécuyer et al.,
1998; Genda and Ikoma, 2008).

As both modelling approach and isotopic measurement of hydrous minerals cannot
provide reliable magnitudes of δ²H change through time, a more appropriate way to decipher
the hydrogen isotope composition of past surface waters is to analyze their fossil remnants
trapped in minerals during their growth. Indeed, surface waters are commonly trapped as tiny
fluid inclusions in the crystal lattice defects of various neoformed minerals. Fluid inclusions,
indeed, represent a unique opportunity for a straightforward determination of some past
climatic parameters (air and water temperatures, air humidity) and the chemical and isotopic
compositions of both ocean and atmosphere. Fluid inclusions were used, for example, to
estimate paleoaltitudes during the evolution of mountain belts (Blisniuk and Stern, 2005;
Mulch and Chamberlain, 2007; Johnston et al., 2018) by analyzing quartz-fill in fractures of
crustal rocks. Fluid inclusions in calcite precipitated as speleothems in caves were also
studied to reconstruct past mean air temperatures (Schwarcz et al. 1976; Harmon et al., 1979; Matthews et al. 2000; Dennis et al. 2001; McGarry et al., 2004; Verheyden et al., 2008; Griffiths et al., 2010; Arienzo et al., 2013; Blamey et al., 2016a; Affolter et al., 2019).
Seawater inclusions in marine carbonates provide an opportunity to analyse chemical variations of both ancient atmosphere and ocean (Johnson and Goldstein, 1993). Residual evaporated waters are also trapped in evaporitic minerals and especially within one of critical interest, NaCl (halite), which is nominally H⁻ and O⁻-free. Therefore, fluid inclusions in sedimentary halite from marine evaporite formations were used to estimate changes in chemical composition of Proterozoic (Kovalevych et al., 2006) and Phanerozoic seawater (Kovalevych et al., 1998, Lowenstein et al., 2001, Horita et al., 2002, Brennan and Lowenstein, 2002, Petrychenko et al., 2005; Timofeeff et al., 2006; Brennan et al., 2013; Meng et al., 2018). More recently, they have been used to quantify the Neoproterozoic molecular oxygen partial pressure of the Earth's atmosphere (Blamey et al., 2016b).
Hydrogen and oxygen isotope compositions of water inclusions in halite were also determined in order to identify fluid sources and mixing processes (Knauth and Beeunas, 1986; Horita 1990; Yang et al. 1995; Rigaudier et al., 2011).

In this study, we measured the deuterium content in water present in tiny primary fluid inclusions trapped in Mesoproterozoic (≈ 1.4 Ga) and Neoproterozoic (≈ 0.8 Ga) halite crystals in order to decipher whether or not the Proterozoic hydrosphere was deuterium-depleted relative to present-day's.

## 2. Material and methods

### 2.1. Sample collection and geological setting

#### 2.1.1. Neoproterozoic Browne Formation, Australia
We studied halite samples from the Lancer 1 and Empress 1A drill cores located close to each other in the northwestern part of the Officer Basin (Figure 1). Halite was sampled in the “B” interval of the Browne Formation whose age is constrained between 800 and 830 Ma on the basis of δ¹³C and ⁸⁷Sr/⁸⁶Sr global correlations, and U/Pb zircon dating of the Gairdner Dyke Swarm (827±6 Ma; Wingate et al., 1998) and the Rook Tuff (802±10 Ma; Fanning et al., 1986). Stevens and Apak (1999) described the sedimentary succession at Empress-1A, it is made of evaporites such as gypsum, anhydrite and halite along with calcareous shales, sandstones, dolomitic limestones and stromatolitic dolomites. Townson (1985) and Carlsen et al. (1999) proposed that this sedimentary sequence was deposited in a shallow marine environment, possibly an interior sea, with prevailing hot and dry climatic conditions. Those marine waters were, however, regularly subject to freshwater excursions to generate brief episodes of brackish water. Townson (1985) also proposed that these sedimentary series reflect the ephemeral existence of coastal lakes, e.g., like playa lakes that occur in the present-day Southern High Plains of the United States or in the Ebro Basin (Salvany et al., 1994).

## 2.1.2. The Mesoproterozoic Sibley Group, Canada

Halite samples were obtained through drilling of the Mesoproterozoic sedimentary rocks of the Sibley Group that crops out between Lake Nipigon and Lake Superior, northern Ontario, Canada (Figure 2). According to Rogala et al. (2007), U-Pb zircon dating indicates a maximum age of deposition of 1443±31 Ma and the sedimentary succession resembles that documented in the Neoproterozoic Browne Formation, Australia. Indeed, the deposits consist of alternating silstones and dolostones, which indicate hydrological cyclic variations of the lacustrine system (Talbot and Allen, 1996). Nodular chert, calcite blebs and the increasing occurrence of evaporites reflect the transition from a lacustrine or deltaic environment to an
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ephemeral playa lake during intense periods of aridity (Cheadle, 1986; Rogala et al., 2007).

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The uppermost lacustrine strata that represent the area that the salt would come from record
an abrupt increase of δ³⁴S from from 5 ‰ to 12 ‰ (CDT) and the Y/Ho ratio increases as
well (Metsaranta, 2006). Both of these geochemical data indicate the possibility of a limited
marine incursion in the upper Channel Island member of the Rossport Formation, which
represents the time interval that the lake shrank and dried up, ie. when the salt was deposited.

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## 2.2. Integrity of halite fluid inclusions

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Only halite samples from the Neoproterozoic Browne Formation and from the Mesoproterozoic Sibley Group, which fulfilled the multi-step screening method developed by Blamey and Brand (2019) for selecting the most pristine halite crystals, were selected for the determination of water δ²H. Detailed methods and results are given in Blamey et al. (2016b), Blamey and Brand (2019) and Blamey et al. (2018). In summary, the halite samples satisfied the following criteria with the associated outputs:

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1) Petrography for studying that support the preservation of primary fluid inclusions, Petrography was used to identify and select sub-samples that only contain primary fluid inclusions based on fabrics and texture such as chevron bands, cornets, and hoppers (see photomicrographs published in Blamey and Brand, 2019). Most fluid inclusions have sizes ranging from 5 µm to 50 µm (see photomicrographs published in Blamey et al., 2016b and Blamey et al., 2018).

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2) Microthermometry to define the range of temperature of crystallization. Homogenization temperatures obtained for the Tonian Browne Formation are comprised between 40°C and 47°C (Blamey and Brand, 2019), while those for the Ectasian-Calymmian Sibley Group range from 22.3°C to 33.7 °C (Blamey et al.,

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191 2018). Such temperatures are compatible with low latitude Earth's surfaces exposed
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193 3) Trace element (Br, Ca, Mg, Sr), and sulfur and strontium isotope geochemistry to check the preservation of seawater chemistry in the case of evaporites of marine origin. Both studied salt deposits suggest the presence of a marine component, more particularly for the Browne Formation (Blamey et al., 2016b).

197 4) Argon isotope geochemistry to track any possible leakage of gas from fluid inclusions.
198 Elevated ⁴⁰Ar/³⁶Ar values relative to modern atmosphere for the ancient samples
199 reflect in situ production of radiogenic argon by the decay of ⁴⁰K housed in fluid
200 inclusions. Data suggest that fluid inclusions behaved as closed systems with respect
201 to argon since their deposit about 1.4 Ga and 0.8 Ga ago, respectively (Blamey et al.,
202 2016b; Blamey and Brand, 2019; Blamey et al., 2018).

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204 2.3. Isotopic analysis of halite fluid inclusions

205 2.3.1. Limit of the method

206 Off-line techniques (Kishima and Sakai, 1980; Vennemann et al., 1993; Lécuyer et al., 1994; Morrison et al., 2001) allow to determine both δ²H and δ¹⁸O ratios (‰ VSMOW) of water inclusions pending that large amounts (several grams) of halite crystals are available per aliquot. Unfortunately, such amount is not available for samples obtained from drilling cores, particularly after a cautious selection of the grains that preserved primary fluid inclusions. Only a few tens of milligrams can be used for the isotopic analysis. Theoretically, both isotopic measurements could be performed on-line using carbon-based reactors that potentially provide oxygen as CO and hydrogen as H₂ for isotopic analysis. However, Gehre et al. (2017), have shown that an accurate hydrogen isotopic analysis of halogen-bearing matrices is precluded because of the "formation of hydrogen-bearing reaction products other
than molecular hydrogen ($H_2$) responsible for non-quantitative $H_2$ yields and possible
hydrogen isotopic fractionation". So far, only Cr-based reactors are able to provide an
accurate $^{2}$H/$^{1}$H determination of water inclusions in halite (Fourel et al., 2019). However, the
measurement of the oxygen isotope ratio is compromised because $H_2$ is obtained by reduction
of $H_2O$ in presence of Cr metal at high temperature within the reactor of the elemental
analyzer, which means that oxygen reacts with Cr to form $Cr_2O_3$.

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223 2.3.2. Instrumental configuration for the $\delta^{2}$H measurement of water inclusions

The protocol used in this study was directly inspired by the one originally proposed by
Fourel et al. (2019). In summary, $\delta^{2}$H measurements of water trapped as inclusions in halite
were performed using a Pyrocube® Elemental Analyser (Elementar GmbH-Germany)
connected online in continuous flow mode to an IsoPrime IRMS. We used a method where
the Pyrocube® elemental analyzer operating at a pyrolysis temperature of 1450°C was
configured with a chromium-based ceramic reactor allowing the $\delta^{2}$H of halite fluid inclusions
to be measured accurately (Fourel et al., 2019). Measurements were calibrated using
reference material (Gerstenberger and Herrmann, 1983; Hut, 1987; Coplen et al., 2006)
IAEA-CH7 ($\delta^{2}$H = -100.3±2.0 ‰ VSMOW), NBS22 ($\delta^{2}$H = -120±1 ‰ VSMOW;
Silverman, 1964; Coplen et al., 2006; Qi et al., 2004), NBS30 ($\delta^{2}$H = -65.7±0.3 ‰
VSMOW; Gonfiantini, 1984), the Pollen working standard ($\delta^{2}$H = -156.1±2.0 ‰ VSMOW,
J. Morrison, personal communication, 2017), L-tryptophane ($\delta^{2}$H = -48±3 ‰ VSMOW, P.
Jame, personal communication, 2017), and L-valine ($\delta^{2}$H = -101±3 ‰ VSMOW, P. Jame,
personal communication, 2017).

Grain size of halite samples is close to 1 mm, large enough to analyze both small and
large fluid inclusions. The use of larger grain size may preclude the deprecitation of all fluid
inclusions. This method has been already tested with success by analyzing 10 mg aliquots of
241 synthetic halite crystals with known δ²H from their fluid inclusions having H concentrations
242 at the 0.2% level (Fourel et al., 2019). The reproducibility associated with these
243 measurements was close to ±3 ‰. In the case of the Proterozoic halite samples, aliquot size
244 ranged from 50 to 80 mg of halite because they contain less water than synthetic halite. For
245 each halite sample, repeated measurements (n = 3 or 4) provided standard deviations (S.D.)
246 from 2.5 ‰ to 8.8 ‰ with an average value of 5.5 ‰ (Table 1), which is higher than those
247 obtained for the reference materials (0.3 < S.D. < 3.0).

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249 3. δ²H and [H] of Precambrian halite fluid inclusions

250 The eight samples of halite from the Mesoproterozoic Sibley Group have δ²H values
251 ranging from -74 ‰ to -54 ‰ (VSMOW) with a mean value of -61.9±17.2‰ and H
252 concentrations comprised between 44 and 186 ppm (mean = 150±92). The six samples from
253 the Neoproterozoic Browne Formation have δ²H values that range from -89 ‰ to -38 ‰ with
254 a mean value of -63.5±7.2‰ and H concentrations comprised between 80 and 324 ppm
255 (mean = 104±51). We observe that there is no significant correlation between the δ²H and H
256 concentrations for samples from either the Sibley Group (R² = 0.11; n = 8) or the Browne
257 Formation (R² = 10⁻⁴; n = 6) (Figure 3).

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259 4. Discussion

260 4.1.1. Extraction of fluid inclusions from halite: validity of the method

261 Depending on the choice of extraction techniques of fluid inclusions from halite, the
262 recovery of water and the integrity of its stable isotope compositions may be a sensitive issue
263 as demonstrated by Horita and Matsuo (1986). Especially in the case of melting techniques or
264 high-temperature decrepitation (the case in this study), the H and O isotope ratios of released
265 fluids can be compromised owing various reactions that may take place at such high-
temerature of 1450°C. There is a safe way, however, to check if the measured δ²H values
reflect those of water inclusions in halite. It consists in analyzing the δ²H of the water
extracted from synthetic halite crystals precipitated in the laboratory and to compare it with
the composition of the mother water from which halite crystallized (Horita and Matsuo, 1986;
Rigaudier et al., 2011; Fourel et al., 2019). Thus, for example, Rigaudier et al. (2011)
observed that the δ²H of evaporating waters sampled between the stage of halite nucleation
and that of final stage of crystallization bracket the compositions of water inclusions in halite.
Indeed, as water inclusions are continuously trapped in the defects of the halite crystal lattice
throughout their growth, they record the isotopic composition of the residual waters across a
time range during which the isotopic values of the water still evolve. This mechanism of
crystal growth and fluid inclusion trapping explains the standard deviation exceeding the
instrumental error we observe for calibrated materials (see section 2.3.3). For example,
crystallization of halite performed outdoor by Fourel et al. (2019) revealed that the mean δ²H
value (+15.8±6.6‰) of halite fluid inclusions lies between the values of evaporating waters at
the beginning (+8.3±0.7‰) and end (+19.5±1.4‰) of halite formation.

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4.1.2. Meaning of the large variations in δ²H

The relatively large δ²H ranges measured either for the Browne Formation or Sibley
Group halite samples could either record a progressive leaking of fluid from the halite
inclusions or the evolution of an evaporating water mass. The absence of correlation between
the δ²H values and the amount of water trapped in halite constitutes a criterion against leaking
fluid inclusions through time. Indeed, fluid inclusion leakage is expected to generate isotopic
fractionation during water loss according to a process of Rayleigh distillation. Therefore, we
consider further in the discussion that the most likely interpretation of such a range in δ²H

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values of halite fluid inclusions results from the progressive evaporation of a water mass
during the nucleation and growth of halite crystals.

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## 4.2. Possible sources of water trapped in Proterozoic halite

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Sedimentary environments of the Neoproterozoic Browne Formation and Mesoproterozoic Sibley Group suggest that surface waters may have been mixtures of seawater and freshwater in unknown proportions at the time of halite deposition. In the case of the 1.4 Ga halite of the Sibley Group, Canada, a fresh water lake environment has been proposed by Rogala et al. (2007). However, the abrupt increase of ³⁴S/³²S ratios up to +12 ‰ recorded at the top of the sequence suggests a possible marine input at the time the salt was deposited. Indeed, such δ³⁴S value is comparable to those measured by Strauss (1993) and Strauss and Veizer (2008) for the same geological period.

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In the case of the 800 Ma halite of the Browne Formation of the Officer Basin located in Western Australia, a marine environment has been proposed on the basis of sulfur and strontium isotope data (Blamey et al., 2016b; Zaky et al., 2019). Indeed, those halites have δ³⁴S of +15‰ that are close to the value of +17‰ proposed by Canfield (2005) for Neoproterozoic seawater. However, as seawater is a highly concentrated aqueous solution relative to fresh waters (rainfall, rivers, lakes), it buffers the isotopic ratios of brackish waters even for mixtures largely dominated by the freshwater end-member. Let us consider a mixture of seawater (sw) and river water (rw) with a [S]sw/[S]rw ratio of 340 (similar to present-days; Brown et al., 1989) and δ³⁴S of +17‰ and 5‰ (Strauss and Veizer, 2008), respectively. Such mixture will have a δ³⁴S of +15‰ even though its seawater fraction will be only 0.02.

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In terms of hydrogen isotope ratios, such mixture will be dominated by the
composition of the freshwater end-member because the hydrogen molar ratio between both

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end-members is one. Consequently, the δ²H of halite water inclusions will be powerful to detect any presence of freshwater during evaporitic stages of the basin. Such isotopic signatures should strongly depend on the latitudes of these deposits considering the known physical processes that trigger the surface water cycle.

Paleomagnetic studies performed by Li and Evans (2011), Wen et al. (2013) and Pisarevsky et al. (2013) allowed these authors to reconstruct paleogeographic maps that locate Australia 800 My ago, and Canada 1.4 Gy ago, within the intertropical belts (Figure 4). Assuming tropical temperatures similar to present-days, those Proterozoic shallow waters submitted to evaporation became most likely deuterium-enriched by about 30 ‰ to 90 ‰ (median = +50 ‰) during the course of halite crystallization. Indeed, such range of deuterium-enrichment was measured by (Rigaudier et al., 2011) who measured Messinian halites from Sicily, Italy, and which precipitated at temperatures close to 35°C according to microthermometric data. Accordingly, it means that the initial δ²H values of those Proterozoic waters were initially comprised between -140 ‰ and -90 ‰ (VSMOW).

Even in the case of the extreme scenario that Proterozoic halite was deposited in freshwater-dominated or pure freshwater environments, such values are much lower than those documented today within the intertropical belt with δ²H values of meteoric waters ranging from -47 ‰ to +8 ‰ along the island coasts and from -45 ‰ to -3 ‰ along the continental coasts according to the IAEA-WMO database (Figure 5). Consequently, on the basis of our current knowledge of the water cycle, such low Proterozoic δ²H values do not appear compatible with the intertropical paleo-position of both studied Mesoproterozoic and Neoproterozoic sedimentary sequences.

One possible interpretation is to consider that the playa lakes were fed by fluvial waters of which the ultimate source would have been located at high altitude. However, today, even Himalayan rivers, which are located between 24°N and 28°N, have δ¹⁸O values
uniformly distributed between -11 ‰ and -6 ‰ (Gajurel et al., 2006), which roughly correspond to δ²H values between -80 ‰ and -40 ‰ according to the Global Meteoric Water Line (Dansgaard, 1964). These isotopic compositions remain lower than those that fed the Proterozoic playa lakes.

We can also consider that the mother water of evaporites was a highly saline groundwater, which means that less evaporation would be necessary for halite precipitation. If such hypothesis cannot be fully excluded, it remains however unlikely as there is no geological evidence for supporting it.

A challenging interpretation is to postulate that Proterozoic meteoric waters derived from a deuterium-depleted oceanic source. Indeed, such meteoric waters would have derived from an oceanic reservoir depleted by several tens of ‰ relative to SMOW, thus implying geological or extra-terrestrial processes responsible for an increasing deuterium-enrichment of seawater through time.

### 4.3. Plate tectonics and the isotopic imbalance between mantle and seawater

First of all, plate tectonics removes “water” from the oceans as OH⁻ during the hydroxylation of mid-ocean ridge basalts. Indeed, significant amounts of water (between 0.4 and 4 times the present mass of the oceans) are stored in the Earth’s mantle in both hydrous and anhydrous minerals (Michael, 1988; Finger et al., 1989; Bell and Rossman, 1992; Thompson, 1992; Langmuir, 1994; Hirschmann, 2006). Most of this water is concentrated in the Mantle Transition Zone dominated by minerals such as wadsleyite and ringwoodite that may contain up to 3 wt% of water (Pearson et al., 2014; Fei et al., 2017). The origin of this water could be a mixture between juvenile and surface water recycled from Earth’s hydrosphere. Indeed, the main metamorphic minerals formed during basalt-seawater hydrothermal interactions are talc, chlorite, serpentine, prehnite and zeolite in replacement of
clinopyroxene, orthopyroxene, olivine and plagioclase. Hydroxylation of nominally-anhydrous minerals (NAM) is characterized by a ²H/¹H fractionation between mineral and water that ranges from -100 to -30 ‰ depending on crystal lattice and chemistry (Suzuoki and Epstein, 1976; Graham et al., 1984). Outgassing water of mantle origin is quite negligible as most water released at island arcs is recycled surface water trapped into the sediments and hydrothermally-altered oceanic crust of the subducting slab (Plank et al., 2013). Water at Mid-Ocean Ridges is not degassed as the pressure is too high at a depth of about 2000 m. The main source of deep water released to the surface comes from subaerial and aerial intraplate volcanoes (e.g., Foulger, 2005; Mazza et al., 2019). Consequently, imbalance between the rate of entrapment of water during the hydrothermal alteration of basalts at oceanic ridges and the rate of release of water from the mantle during the dehydration of subducting plates is the most likely scenario to happen.

The geological record provides conflicting interpretations, which is most likely due to an indirect estimation of seawater δ²H through the measurement of rock-forming hydrous minerals (antigorite, lizardite, chrysotile, chlorite) as well as full or partial isotopic re-equilibration during geological events (burial, contact metamorphism, volcanic activity and orogenesis) that involve waters of magmatic, metamorphic or meteoric origin (Sheppard, 1986). A synthetic overview of available data in the literature helps to figure out this key issue.

For example, Taylor and Magaritz (1975) and Taylor (1977) noted that the hydrogen isotope composition of gneisses and granites from the Archean Swaziland cratons are about 30 ‰ lower than those belonging to Phanerozoic batholiths. In addition, Taylor (1977) proposed that these low δ²H values could be linked to the hydrogen isotope composition of the altered oceanic crust that may have reacted with an ocean whose δ²H value was closer to -30 ‰ than 0 ‰. In a similar way, we could consider that the low δ²H values measured in
Archean greenstone belts are also compatible with an Archean hydrosphere that was
deuterium-depleted in comparison to modern times. Indeed, Pope (2011) measured δ²H
values of Isua (Greenland) serpentines in the range -99 ‰ to -53 ‰. Lécuyer et al. (1994)
reported amphibole values from -88 to -77 ‰ for the 3.5 Ga old altered komatiites from the
Barberton Greenstone Belt, South Africa, whereas Kyser et al. (1999) measured δ²H of
serpentines from Abitibi, Canada, with a large range from -147 to -54 ‰. It is worth noting
that δ²H values of whole rock samples from a hydrothermally altered layered flow from the
Chukotat Group of the Lower Proterozoic (2.0-1.9 Ga) Cape Smith fold belt (northern
Quebec) range from -53 ‰ to -40 ‰ (Lécuyer et al., 1996), only matching the highest δ²H
values documented (-80 to -30 ‰) in the modern oceanic crust (Sakai et al., 1990; Agrinier et
al., 1995; Früh-Green et al., 1996; Alt and Shanks, 2006. Moreover, many Phanerozoic
ophiolites display δ²H values of serpentine in a global range of -190 to -60 ‰ (Wenner and
Taylor, 1973; Magaritz and Taylor, 1974; Burkhard and O'Neil, 1988; Yui et al., 1990;
Liakhovitch et al., 2006). Consequently, it appears that the Precambrian record matches the
Phanerozoic one with huge isotopic ranges that preclude either any identification of secular
trend or any conclusion regarding the hydrogen isotope composition of Precambrian surface
waters.

407

4.4. Secular unidirectional water fluxes

4.4.1. Photodissociation of water in the upper atmosphere

The present-day flux of hydrogen escape in the upper atmosphere is about 3x10⁸ atom H.cm⁻².s⁻¹ (Yung et al., 1989). This flux is three orders of magnitude lower than water fluxes affecting the mantle-ocean system and thus cannot significantly affect the mass of the hydrosphere. Assuming that hydrogen is entirely lost as ¹H with no loss of ²H, the change in the δ²H value of the oceans was only +2 to +3 ‰ when integrated over a period of 4.5 Gyr

410

411

412

413

414
415 (Yung et al., 1989; Lécuyer et al., 1998). However, it is unlikely that this flux was constant
416 over time, especially during the Precambrian when the Earth's atmospheric chemistry
417 strongly differed in terms of CO₂, O₂ and H₂ partial pressures. Indeed, Holland (1984)
418 considered that the input H flux to the early Earth was much higher, however, in a poorly-
419 constrained range of 1.9x10⁹ to 1.9x10¹³ cm⁻².s⁻², partly resulting from the volcanic activity
420 releasing reduced gases to the atmosphere. Yung et al. (1989) calculated a flux averaged over
421 the Earth's history of 1.3 to 5.8x10¹⁰ cm⁻².s⁻² lying within the range estimated by Holland
422 (1984). For that purpose, Yung et al. (1989) considered that the ²H/¹H ratio of a deep-sea
423 tholeiitic glass, which trapped and preserved primordial rare gases such as He and Ne (Craig
424 and Lupton, 1976), had a hydrogen isotope ratio reflecting that of primordial water.
425 Following this hypothesis, the hydrogen isotope composition of primordial oceans should
426 have been deuterium-depleted by about -80 ‰ relative to SMOW, thus corresponding to a
427 water loss of about 36 %, and strongly questioning the paradigm of the continental freeboard
428 constancy (e.g., Eriksson, 1999). A deuterium enrichment of the hydrosphere may also have
429 been triggered by a longer residence time of H₂ in an oxygen-poor Proterozoic atmosphere.

430 However, it mainly concerns the Mesoproterozoic with a pO₂ estimated close to 8% to 10%
431 Present Atmopsheric Level (Blamey et al., 2019) while the oxygenation level of the
432 Neoproterozoic atmosphere was high enough, with a PAL value close to 50-60%, to preclude
433 a large atmospheric residence time of molecular hydrogen (Blamey et al., 2016b; Blamey et
434 al., 2019).

435

436

4.4.2. Extra-terrestrial addition of water to the Earth
437

Lécuyer et al. (1998) calculated that the total water mass added to the terrestrial oceans in 4.5 Gyr should lie between 4 and 22 % according to the relative contribution of comets and water-rich carbonaceous chondrites. Considering a maximum possible δ²H value for this extraterrestrial water of +990±160 ‰, which matches the P/Halley comet composition (Deloule and Robert, 1995; Eberhardt et al., 1995), the δ²H values of the global ocean should have been close to -10 ‰ (SMOW) 4.2 Gyr ago.

438

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440

441

442

443

444

### 4.4.3. Continental growth and water storage

445

446

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### 4.4.4. H₂O reduction into H₂ during serpentinization of mantle rocks

457

458

459

460

The sum of meteoric waters stored in the atmosphere, lakes, rivers, soils, and groundwaters is about 3.5 % of the total Earth's surface water (Oki, 2006). The δ²H value of meteoric waters from the equatorial-tropical to polar areas range from +30 to -260 ‰ but most of the precipitations fall over equatorial and temperate belts, therefore we can estimate an average δ²H value of -55 ‰ (Lécuyer, 2013). The contribution of this reservoir to the hydrosphere is thus estimated to be close to 2 ‰ and can be reasonably ignored in the mass balance calculation. Lécuyer et al. (1998) calculated a δ²H value for a continent-free Earth's ocean of -18(±6) ‰ and a mass of 1.7(±0.2)×10²¹ kg, about 20 % larger than the present-day mass (1.38×10²¹ kg) of the oceans. This isotopic difference of about -20‰ corresponds to the amount of water stored in the biosphere, the continents, and ice sheets of the modern Earth.
magnitude isotopic fractionation that takes place between H₂O and H₂ during the oxidation of iron-bearing silicates (Simon et al., 2011). As temperatures of serpentinization do not exceed 400°C (Bach et al., 2004; Lamadrid et al., 2017), we calculate 1000lnα(H₂O–H₂) values that increase from about 200 to 450 for temperatures decreasing from 400°C to 100°C (Figure 6) using the isotopic fractionation equation of Simon et al. (2011).

The global annual molar flux of H₂ released to the atmosphere is the subject of various estimates in the literature as this one must also include the serpentinization of both continental and oceanic ultramafic rocks. Rona et al. (2013) estimated a flux at mid-ocean ridges of 1.67x10¹¹ mol H₂·yr⁻¹ whereas Lollar et al. (2014) proposed a global flux in the order 10¹⁰ to 10¹¹ mol H₂·yr⁻¹. Several authors proposed a global H₂ flux close to 10¹¹ mol H₂·yr⁻¹ (Canfield et al., 2006; Emmanuel and Ague, 2007; Sleep and Bird, 2007), but higher estimates close to 10¹² mol H₂·yr⁻¹ were recently proposed by Worman et al. (2016). Andreani et al. (2017) argued on the basis of data and modelling that this flux was not constant through time, but started at 160 Ma with “elevated H₂ rates (>10¹² to 10¹³ mol·yr⁻¹) compared to Late Mesozoic (<160 Ma) rates (<10¹¹ – 10¹² mol·yr⁻¹)”. Here, we propose a simple mass balance calculation based on a Rayleigh-type distillation equation for estimating the magnitude of deuterium-enrichment of the global ocean over 4.5 Gyr. Accordingly, for a temperature of serpentinization ranging from 150 to 350°C, a deuterium-enrichment in the order of a few tens of permil (≤ 80 ‰) is thus expected for fluxes comprised between 10¹¹ and 10¹² mol H₂·yr⁻¹ (Figure 7).

481

482 4.5. Cumulative effects and the δ²H of Proterozoic seawater

483
484

The anomalously low δ²H values obtained for the water trapped into the studied Proterozoic halites could result from the cumulative effects produced by the physico-chemical processes described above:

485

486

487

- addition of extraterrestrial water seems to be negligible in the total isotopic budget of the global ocean

488

489

- continental growth only had a restricted impact on the δ²H of seawater during the Archean as the major part of continents were already formed during the Proterozoic (e.g., Dhuime et al., 2012)

490

491

- the absence of ice sheets contributes about -10‰

492

- photodissociation of water in the upper atmosphere and plate tectonics contributed no more than -20 ‰ according to model outputs about 1.5 Gy ago (Lécuyer et al., 1998; Kurokawa et al., 2018). However, Yung et al. (1989) proposed a deuterium-enrichment up to 80 ‰ by considering higher H₂ fluxes during the photodissociation of water during the Archean.

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497

- another significant contribution (up to 80‰) could be the reduction of water into H₂ during the serpentinization of mantle peridotites. It would imply a H₂ flux ranging from 10¹¹ to 10¹² mol.yr⁻¹ in agreement with some recent published estimates.

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503

## 5. Conclusions

504

505

Well-preserved halite crystals from the Neoproterozoic (800 Ma) Browne Formation,

506

Australia, and Mesoproterozoic (1.4 Ga) Sibley Group, Canada, contain water with low δ²H values ranging from -89 ‰ to -38 ‰ (VSMOW) that cannot be explained in the light of our knowledge of the modern global water cycle. As water strongly evaporates during the course

507

508
509 of halite crystallization, the initial δ²H of water are estimated to be between -140 ‰ and -90 ‰ (VSMOW). Such values are unlikely considering the tropical positions occupied by both
510 basins during the Proterozoic, however, they could testify of the existence of deuterium-
511 depleted oceans compared to those during Phanerozoic times. Physico-chemical processes
512 potentially responsible for a progressive deuterium-enrichment over time have been reviewed,
513 including water flux imbalance between mantle outgassing and water carried by subducted
514 plates, addition of extraterrestrial water and photodissociation of water in the upper
515 atmosphere. However, the isotopic fractionation of large magnitude that takes place during
516 H₂O reduction into H₂ is a key chemical reaction that accompanies the serpentinization of
517 ultramafic rocks, which could represent the main process responsible for the progressive
518 deuterium-enrichment of the oceans by a few tens of per mil.
519

520

521 Acknowledgements – This study has been founded by the Tellus-SYSTER CNRS Program
522 2018 (CL). The authors also thank one anonymous reviewer and W. M. White for their
523 valuable and constructive reviews that helped us to improve the scientific content of this
524 study.

525
526

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878

879 Table captions:

880

881 Table 1: δ²H (‰ VSMOW) and [H] concentrations of fluid inclusions trapped in the studied

882 Proterozoic halite crystals. Samples L (Lancer1) and E (Empress1A) come from the Officer

883 Basin (Figure 1), Australia (800 Ma), and samples P1 to P8 come from the Mesoproterozoic

884 Sibley Group (Figure 2) (1.4 Ga). S.D. = standard deviation. N = number of measurements.

885

886 Figure captions:

887

888 Figure 1: A) Geographic map of Australia with the location of Officer Basin and the two

889 Lancer 1 and Empress 1A drill cores. B) Sampling horizons in the Browne Formation.

890 Numbers indicate the position of halites that have been measured for the δ²H (‰ VSMOW)

891 of their water inclusions (Table 1).

892

893 Figure 2: Geographic map of Canada (A) with the location of Thunder Bay (black and red

894 dot), Lake Superior (B), and the outcrop of Sibley Group (C). Samples P1 to P8 belonging to

895 the Outan Island Formation (D), Sibley Group, have been measured for the δ²H (‰

896 VSMOW) of their water inclusions (Table 1).

897

898 Figure 3: Variations of δ²H of water inclusions in Proterozoic halites as a function as a

899 function of the amount of hydrogen extracted from the sample by pyrolysis.

900
Figure 4: Various scenarios of Earth paleogeographic reconstructions for the A) Calymnian-Ectasian transition (1.4 Ga) and the B) Tonian (≈ 750 Ma) according to Li and Evans (2011), Wen et al. (2013) and Pisarevsky et al. (2013). Note that for any reconstruction Laurentia and Australia continents are located within the intertropical belt.

Figure 5: Frequency histogram of present-day δ²H of meteoric waters from tropical islands and continental coasts. Data were extracted from the IAEA-WMO database.

Figure 6: ²H/¹H fractionation between water and dihydrogen during the oxidation of hydrous iron-bearing silicates. The curve was computed according to the experimental study performed by Simon et al. (2011) who determined the following equation: α(H₂O-H₂) = 0.2447729(±0.027705)·10⁶·T⁻² + 1.024(±0.022).

Figure 7: Deuterium-enrichment of the oceanic reservoir as a function of the H₂ flux released during the serpentinization of continental and oceanic ultramafic rocks. During serpentinization, the oxidation of ferrous iron into ferric iron is accompanied by the reduction of water into dihydrogen. Curves have been computed for three temperatures (T = 150°C with 1000lnα(H₂O-H₂) = 2.4085; T = 250°C with 1000lnα(H₂O-H₂) = 1.9297; T = 350°C with 1000lnα(H₂O-H₂) = 1.6623; see Figure 6) using a Rayleigh distillation equation: R = R₀f(α⁻¹) with R₀ = isotopic ratio of the initial oceanic mass (1.38x10⁻²¹ kg), R = isotopic ratio of the decreasing oceanic mass, f = fraction of the residual oceanic mass, and α = isotopic fractionation between H₂O and H₂. Fluxes ranging from 10¹¹ to 10¹² H₂ mol·y⁻¹ have been
923 integrated over 4.5 Gyr and correspond to a loss of the oceanic mass ranging from 0.6% to

924 5.5%.
<figure><img src="image_2.png" /></figure>

A)

<figure><img src="image_3.png" /></figure>

B)
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B

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<figure><img src="image_6.png" /></figure>

Osler Group

Sibley Group

Diabase

Animikie Group

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D
<figure><img src="image_8.png" /></figure>

$\delta^{2}\text{H}$ of water inclusions in Proterozoic halite
A)

<figure><img src="image_9.png" /></figure>

B)

<figure><img src="image_10.png" /></figure>
<figure><img src="image_11.png" /></figure>
<figure><img src="image_12.png" /></figure>
<figure><img src="image_13.png" /></figure>
<table><thead><tr><th>Sample</th><th>δ<sup>2</sup>H</th><th>S.D.</th><th>N</th><th>H (ppm)</th></tr></thead><tbody><tr><td colspan="5"><em>Officer Basin, Australia</em></td></tr><tr><td>L1466</td><td>-60.4</td><td>6.3</td><td>3</td><td>93</td></tr><tr><td>E1478</td><td>-37.6</td><td>8.5</td><td>4</td><td>80</td></tr><tr><td>E1482</td><td>-70.9</td><td>8.8</td><td>3</td><td>95</td></tr><tr><td>E1492</td><td>-58.3</td><td>6.1</td><td>3</td><td>324</td></tr><tr><td>E1497</td><td>-89.1</td><td>4.6</td><td>4</td><td>130</td></tr><tr><td>E1502</td><td>-54.9</td><td>5.3</td><td>4</td><td>179</td></tr><tr><td colspan="5"><em>Sibley Group, Canada</em></td></tr><tr><td>P1</td><td>-66.2</td><td>3.2</td><td>4</td><td>125</td></tr><tr><td>P2</td><td>-72.5</td><td>5.8</td><td>3</td><td>44</td></tr><tr><td>P3</td><td>-57.9</td><td>3.6</td><td>4</td><td>49</td></tr><tr><td>P4</td><td>-74.1</td><td>3.0</td><td>4</td><td>122</td></tr><tr><td>P5</td><td>-59.6</td><td>7.5</td><td>4</td><td>65</td></tr><tr><td>P6</td><td>-59.1</td><td>6.7</td><td>4</td><td>150</td></tr><tr><td>P7</td><td>-64.9</td><td>5.1</td><td>3</td><td>91</td></tr><tr><td>P8</td><td>-53.8</td><td>2.5</td><td>3</td><td>186</td></tr></tbody></table>