Open Access

Reactions between komatiite and CO2-rich seawater at 250 and 350 °C, 500 bars: implications for hydrogen generation in the Hadean seafloor hydrothermal system

  • Hisahiro Ueda1, 2Email author,
  • Takazo Shibuya2, 3, 4,
  • Yusuke Sawaki1,
  • Masafumi Saitoh2, 3, 4,
  • Ken Takai2, 3, 4, 5 and
  • Shigenori Maruyama6
Progress in Earth and Planetary Science20163:35

DOI: 10.1186/s40645-016-0111-8

Received: 22 March 2016

Accepted: 21 October 2016

Published: 18 November 2016

Abstract

To understand the chemical nature of hydrothermal fluids in the komatiite-hosted seafloor hydrothermal system in the Hadean, we conducted two hydrothermal serpentinization experiments involving synthetic komatiite and a CO2-rich acidic NaCl fluid at 250 and 350 °C, 500 bars. During the experiments, the komatiites were strongly carbonated to yield iron-rich dolomite (3–9 wt.% FeO) at 250 °C and calcite (<0.8 wt.% FeO) at 350 °C, respectively. The carbonation of komatiites suppressed H2 generation in the fluids. The steady-state H2 concentrations in the fluid were approximately 0.024 and 2.9 mmol/kg at 250 and 350 °C, respectively. This correlation between the Fe content in carbonate mineral and the H2 concentration in the fluid suggests that the incorporation of ferrous iron into the carbonate mineral probably limited magnetite formation and consequent generation of hydrogen during the serpentinization of komatiites. The H2 concentration of the fluid at 350 °C corresponds to that of modern H2-rich seafloor hydrothermal systems, such as the Kairei hydrothermal field, where hydrogenotrophic methanogens dominate in the prosperous microbial ecosystem. Accordingly, the high-temperature serpentinization of komatiite would provide the H2-rich hydrothermal environments that were necessary for the emergence and early evolution of life in the Hadean ocean. In contrast, H2-rich fluids may not have been generated by serpentinization at temperatures below 250 °C because carbonate minerals become more stable with decreasing temperature in the komatiite-H2O-CO2 system.

Keywords

Komatiite CO2-rich condition Early Earth Hydrothermal alteration Serpentinization

Introduction

Deep-sea hydrothermal environments have been considered as a favorable environment for the emergence and early evolution of life on Earth (e.g., Yanagawa and Kojima 1985; Russell and Hall 1997). In particular, H2-rich hydrothermal fluids generated through the serpentinization of ultramafic rocks would have driven prebiotic chemical evolution and the development of biotic energy metabolisms (Takai et al. 2006; Amend and McCollom 2009; Russell et al. 2014; Nakamura and Takai 2014; Shibuya et al. 2016). Molecular hydrogen (H2) generation during serpentinization is caused by the reduction of water in conjunction with the oxidation of ferrous iron in silicates. This process is written as the following simplified reaction (e.g., Allen and Seyfried 2003; Seyfried et al. 2007; McCollom and Bach 2009):
$$ 2{\left(\mathrm{F}\mathrm{e}\mathrm{O}\right)}_{\mathrm{rock}} + {\mathrm{H}}_2\mathrm{O}\to {\left({\mathrm{Fe}}_2{\mathrm{O}}_3\right)}_{\mathrm{rock}} + {\mathrm{H}}_2, $$
where the ferric iron-bearing solid phase generally precipitates as magnetite (Fe3O4). Therefore, H2 generation is strongly related to magnetite formation during serpentinization (McCollom and Bach 2009).

Among the various prospects for the seafloor hydrothermal systems in the early Earth (e.g., Kump and Seyfried 2005; Shibuya et al. 2010), two types of serpentinization are hypothesized to be dominant in the Hadean ocean: a komatiite-hosted high-temperature type at oceanic islands/plateaus (Takai et al. 2006; Yoshizaki et al. 2009; Shibuya et al. 2015) and a low-temperature alkaline type hosted by komatiites or peridotites at off-ridge ocean floors (e.g., Russell et al. 2010, 2014; Shibuya et al. 2016). Based on the geological records, Archean oceanic crust was likely much thicker than modern equivalents owing to the higher potential mantle temperature at that time (Ohta et al. 1996; Komiya et al. 2002; Komiya 2004; Moores 2002; Shibuya et al. 2012). This thick lid of oceanic crust probably limited the exposure of mantle peridotites on the seafloor, thus suggesting that komatiite volcanism would have been much more abundant than the exposed mantle peridotites that are frequently observed near modern slow-spreading ridges without a sufficient magmatic supply (Takai et al. 2006). Therefore, it has been pointed out that H2-rich seafloor hydrothermal environments would have been predominantly driven by komatiite volcanism in the Hadean ocean (Takai et al. 2006). On the other hand, the low-temperature alkaline type could provide distinctive chemical environments such as large pH gradients in the seawater-hydrothermal fluid mixing zones, which may also have been advantageous for the possible development of proton-motive energy metabolisms (Russell et al. 2014; Shibuya et al. 2016).

Previously, hydrothermal alteration experiments have been conducted to understand the potential of H2 generation during the serpentinization of komatiites under CO2-free conditions (Yoshizaki et al. 2009; Shibuya et al. 2015; Suzuki et al. 2015). However, as suggested by theoretical calculations (Walker 1985; Kasting 1993; Sleep and Zahnle 2001) and geological records (Lowe and Tice 2004; Ohmoto et al. 2004; Shibuya et al. 2007, 2012, 2013a), the atmospheric CO2 levels in the early Earth were likely much higher than the modern level. Furthermore, it has been experimentally demonstrated that carbonate formation during the serpentinization of olivine under CO2-rich conditions suppresses H2 generation in fluids (Jones et al. 2010; Klein and McCollom 2013; Neubeck et al. 2014). Recently, some experiments have been conducted in komatiite-CO2-H2O systems (Lazar et al. 2012; Hao and Li 2015), but H2 generation in fluids was not the objective in these studies. Therefore, the potential for hydrogen generation through the serpentinization of komatiites in the Hadean has not yet been experimentally evaluated under CO2-rich conditions.

In this study, we conducted two experiments that simulate the reactions between komatiite and CO2-rich seawater at 250 and 350 °C, 500 bars, using a batch-type (closed system) hydrothermal reactor (Yoshizaki et al. 2009). The experiments revealed the chemical composition of high-temperature hydrothermal fluids and CO2 absorption ability of komatiite through serpentinization under CO2-rich conditions. The results demonstrate the significance of komatiite-hosted hydrothermal systems for seawater chemistry and provide insight into the H2-rich hydrothermal environments in the CO2-rich Hadean ocean.

Methods/Experimental

Synthesis of komatiite

The komatiite used in the experiments was synthesized from a mixture of 12 reagents (SiO2, TiO2, Al2O3, Fe2O3, MnO, MgO, CaCO3, Na2CO3, K2CO3, P2O5, NiO, and Cr2O3). The chemical composition of the mixed reagents was adjusted to Al-depleted (Barberton-type) komatiite (Smith et al. 1980; Wei et al. 1990; Yoshizaki et al. 2009) because the volcanism of Al-depleted komatiite was likely more predominant than Al-undepleted komatiite in the Hadean (Shibuya et al. 2015). The mixed powder was placed in a Pt-Rh crucible at 1000 °C for 1 h in an electronic furnace to decarbonate the reagents. The mixture was melted at 1600 °C for 0.5 h at the oxygen fugacity of quartz-fayalite-magnetite (QFM) buffer regulated by a H2-CO2 gas mixture (Canil 1997). To create a spinifex texture of olivine, the temperature was lowered to 1450 °C over 1.5 h. Next, the crucible was quenched to room temperature, thus yielding a fresh spinifex-textured komatiite (Fig. 1). The composition and texture of the synthetic komatiite and its mineral/glass phases were analyzed using X-ray diffraction (XRD), X-ray fluorescence (XRF), and an electron probe microanalyzer (EPMA) (Table 1). In the interstitial glass, dendritic olivine crystals configurate a spinifex texture. Broadly depending on size, these crystals have Mg# values ranging from 89 to 94. Small amounts of chromian spinels are disseminated in the vicinity of the olivine crystals. Owing to such a mineral assemblage, almost all Ca in the komatiite was partitioned into the interstitial glass (Fig. 1 and Table 1). The synthesized komatiite was crushed in a tungsten mortar and sieved to obtain a <100 μm powder. To remove any contamination of organic materials during sample preparation, the powdered komatiite was washed with acetone and distilled water several times and dried for 12 h in an oven at 50 °C.
Fig. 1

a Photomicrograph taken in cross-polarized light and b a backscattered electron (BSE) image of the synthetic komatiite. Ol olivine, Spl spinel, Gl glass

Table 1

Compositions of the synthetic komatiite, mineral/glass phases therein, starting materials used in a previous experiment (wt.%)

Sample

Al-depleted komatiite

Olivine

Glass

Spinel

Basaltb

SiO2

46.29

41.02

50.74

0.22

49.84

TiO

0.37

0.00

0.68

0.22

0.73

Al2O3

4.30

0.20

7.74

6.43

15.82

Cr2O3

0.21

0.48

0.52

52.22

FeOtotal a

12.52

9.70

14.81

22.03

11.00

MnO

0.13

0.09

0.18

0.14

0.22

MgO

27.47

48.80

12.57

15.49

9.44

CaO

7.20

0.30

11.84

0.22

10.84

Na2O

0.06

0.00

0.23

0.00

1.88

K2O

0.02

0.02

0.05

0.02

0.18

NiO

0.11

0.34

0.01

0.22

Total

98.68

100.95

99.37

97.23

99.95

Mg#c

79.64

89.96

60.21

55.62

60.47

–, no mention

aTotal iron as FeO

bSynthetic basalt used in Shibuya et al. (2013b)

cMg# = [Mg / (Mg+Fe)] × 100

Experimental system

The Inconel alloy autoclave used for the hydrothermal alteration experiment is based on the study by Seyfried et al. (1979). It resists corrosion and withstands high-temperature and high-pressure conditions (up to 600 °C and 600 bars; Fig. 2). The reaction cell is made of a gold bag with a titanium head because of the resistance of these materials to high-temperature fluids. Furthermore, gold has sufficient flexibility to allow the fluid inside the reaction cell to be pressurized by the surrounding water. In addition, to avoid possible H2 generation via the reaction of metallic titanium with water, the surface of the titanium was completely oxidized before the experiments. To eliminate organic contamination, all materials in contact with the reacting fluid were baked in a muffle furnace at 450 °C for 5 h.
Fig. 2

Schematic illustration of the Inconel alloy autoclave used in this study (modified after Yoshizaki et al. 2009). The synthesized komatiite and hypothetical Hadean seawater were sealed in the gold bag

The initial solution used in the experiments was controlled at pH 4.9 and 400 mmol/kg of ΣCO2 (=CO2 (aq) + HCO3  + CO3 2−) by the addition of NaHCO3, NaCl, and HCl to ultrapure water, according to the previous studies suggesting that Hadean seawater was acidic due to the high partial pressure of CO2 in the atmosphere (e.g., Kasting 1993; Macleod et al. 1994). To avoid the degassing of CO2 by a reaction between NaHCO3 and HCl prior to the sealing of the reaction cell, these materials were separately placed in the gold tube and then mixed after sealing (Shibuya et al. 2013b). The Cl concentration of the solution was regulated to be approximately 1000 mmol/kg because salinity of seawater in the early Earth was potentially 1.5–2 times higher than the modern value (approximately 550 mmol/kg) because of the absence of continents and the associated salt deposits/saline water (Knauth 2005). The komatiite powder was reacted with this hypothetical seawater in the reaction cell at 250 and 350 °C, 500 bars for 1530–2760 h (hereafter abbreviated as Exp-250 and Exp-350, respectively). At the beginning of the experiments, the water/rock mass ratios were adjusted to approximately 5 because water/rock ratios are commonly below 5 in the high-temperature regions of natural subseafloor hydrothermal systems (Wetzel and Shock 2000). The water/rock mass ratio decreased to ~3 at the end of the experiments due to the multiple fluid samplings during the experiments.

Sampling and analytical methods

During the experiment, approximately 3–4 g of fluid samples was collected several times through a gold-lined sampling tube. For CO2 and H2 analyses, 0.5 mL of fluid sample was directly introduced to each Ar-purged, sealed vial at room temperature. Especially, for the CO2 analysis, the sampled fluid was acidified (pH <2) by adding HCl to ensure the complete extraction of the dissolved bicarbonate and carbonate ions. Quantitative analysis for gas species was conducted by gas chromatography. The overall detection precision for the analyses of ΣCO2 and H2 concentrations in the fluid were both better than 5%. For the analysis of other dissolved species, 0.2 mL of the fluid was collected in two vials, to which either HNO3 or NaOH was immediately added to avoid the precipitation of transition metals (Fe and Mn) and silicic acid species, respectively. The fluid samples were analyzed using inductively coupled plasma optical emission spectrometry and ion chromatography. The analytical precision (2σ) was approximately 2% for Cl and Na, and 5% for the other elements.

The pH of the fluid samples was determined using a pH meter at 25 °C under atmospheric conditions. The significant digits and precision of the pH measurement are 2 digits and ±0.1 units. This measurement was performed 1 h after sampling to allow the stabilization of the pH against CO2 degassing. The pH25 °C of the fluid cannot be measured directly because it rapidly changes due to CO2 degassing after the sampling of fluid from the reactor. The amount of degassed CO2 was determined from the ΣCO2 concentrations of the fluid sample directly introduced into the sealed vial and of the fluid degassed under atmospheric conditions, then the original pH25 °C was calculated based on the amount of degassed CO2. We calculated pHin-situ of the high-temperature fluid in the reaction cell with the Geochemist’s Workbench computer code (Bethke 2008) based on the pH at room temperature (pH25 °C) and concentrations of dissolved elements/species and gases. In the pHin-situ calculations, charge balance was constrained by pH25 °C, while Na was used to compensate for imbalanced charges derived from analytical errors. The required thermodynamic database was generated by the SUPCRT92 computer program (Johnson et al. 1992), using thermodynamic data of minerals, aqueous species, and complexes reported in Shock and Helgeson (1988), Shock and Koretsky (1995), Shock et al. (1989, 1997), Sverjensky et al. (1997), and McCollom and Shock (1997). The B-dot activity model was used in the calculations (Helgeson 1969; Helgeson and Kirkham 1974). The temperature-dependent activity coefficient for aqueous CO2 was derived from the empirical relationship established by Drummond (1981), and the temperature-dependent activity of water in NaCl solution was derived from the formulation of Bethke (2008). Cleverley and Bastrakov (2005) provided useful temperature-dependent polynomial functions for both of these last-mentioned parameters.

After the experiments, alteration products were retrieved from the reaction cell and dried at 80 °C for 12 h. The solid alteration product of each experiment was analyzed with XRD and EPMA to determine the assemblage and composition of the alteration minerals. Powdery alteration products from the experiments at both temperatures were embedded in Epofix resin, and thin sections were made of these. The analytical conditions of the EPMA were 15 kV of accelerating voltage, 10 nA of specimen current, and 60–80 s of counting time. The precision of the elemental concentration was less than 10%, which was based on duplicate measurements. The detection limit for the elemental concentration was 0.01 wt.%. Moreover, the starting solid material (unaltered komatiite) was analyzed with XRD, XRF, and EPMA.

Results and Discussion

Fluid chemistry

The Cl concentration in the fluid was kept relatively constant throughout both experiments because Cl-bearing minerals are generally rare in hydrothermally altered rocks (Fig. 3 and Table 2). In contrast, the concentrations of all other components (Na, K, Mg, Ca, Si, Fe, Mn, and ΣCO2) changed immediately after the beginning of the experiments and reached a near steady state within 2760 h at 250 °C and 1530 h at 350 °C. The Mg concentration in Exp-250 (36–40 mmol/kg) was 30–40 times higher than that in Exp-350 (<1.2 mmol/kg). The Fe concentration in Exp-250 (0.49–0.97 mmol/kg) was approximately half than that in Exp-350 (1.53–1.80 mmol/kg). No large difference between Exp-250 and Exp-350 was identified in the concentrations of other dissolved elements. Na (820–1036 mmol/kg across both experiments) had the second highest concentration next to Cl. The concentration of Ca increased and converged at 39–49 mmol/kg in Exp-250 and 37–42 mmol/kg in Exp-350. The concentrations of Si and K were 1.6–6.1 and 0.78–1.24 mmol/kg, respectively, across both experiments. The concentration of Mn was the lowest among all analyzed elements (0.15–0.73 mmol/kg across both experiments).
Fig. 3

Concentrations of the dissolved species in an aqueous fluid coexisting with komatiite powder and its alteration products as a function of reaction time in the experiments at a 250 and b 350 °C, 500 bars

Table 2

Composition of the sampled fluids in the experiments (mmol/kg)

Experiment

Time (h)

pH25 °C a

ΣCO2 b

pH25 °C c

pHin-situ d

H2

ΣCO2

Cl

Na

K

Mg

Ca

Si

Fe

Mn

250 °C

0

n.a.

n.a.

4.9

6.0

0.003

396

1128

1036

0.08

n.d.

0.073

n.d.

n.d.

n.d.

43

6.3

0.9

3.7

4.1

0.010

161

1174

939

0.94

36.2

46.5

6.1

0.49

0.19

119

5.8

1.9

4.1

4.5

0.014

67

1168

951

0.89

40.4

39.7

4.4

0.56

0.15

452

6.6

1.0

4.3

4.7

0.004

35

1186

948

0.85

38.4

39.2

2.4

0.77

0.23

859

6.4

2.3

4.6

5.0

0.007

36

1178

937

1.24

37.2

48.7

2.1

0.85

0.24

1771

7.1

1.0

4.4

4.8

0.021

35

1186

820

0.88

36.4

43.6

1.6

0.93

0.24

2760

6.8

1.2

4.5

4.8

0.024

33

1175

949

0.95

36.4

48.8

3.5

0.97

0.25

350 °C

0

n.a.

n.a.

4.9

7.2

n.d.

396

966

946

0.10

0.03

0.078

n.d.

n.d.

n.d.

90

6.1

1.9

3.8

5.6

6.0

176

971

902

0.78

1.2

39.4

3.7

1.53

0.71

258

6.4

1.7

3.9

5.6

2.8

163

968

882

0.86

1.2

38.0

3.7

1.80

0.73

594

6.4

1.7

3.9

5.6

2.9

179

963

877

0.98

1.1

42.2

3.6

1.65

0.67

1146

6.5

1.8

3.9

5.7

2.7

174

969

900

1.08

1.0

41.1

3.8

1.61

0.63

1530

6.6

1.7

3.9

5.7

2.2

171

953

884

0.81

1.0

37.3

4.1

1.55

0.60

n.a., not analyzed

n.d., not detected

aMeasured pH after CO2 degassing

bMeasured ΣCO2 concentration after CO2 degassing

cCalculated pH25 °C at 1 bar before CO2 degassing at room temperature

dCalculated pHin-situ at high-temperatures, 500 bars

The ΣCO2 concentration decreased from ca. 400 mmol/kg at the beginning of the experiments to approximately 33 mmol/kg in Exp-250 and approximately 171 mmol/kg in Exp-350 (Fig. 4a and Table 2). This trend of the steady-state ΣCO2 concentration in the higher-temperature experiment being higher than that in the lower-temperature experiment is consistent with the results of experiments simulating high-temperature reactions between basalt and CO2-rich seawater (Shibuya et al. 2013b). On the other hand, the pHin-situ values were higher than the pH25 °C values at the beginning of the experiments due to the temperature dependence of the ion product of water and speciation of CO2 species (e.g., Shibuya et al. 2010). As the reaction proceeded, the pHin-situ values finally decreased to 4.8 in Exp-250 and 5.7 in Exp-350, which are broadly consistent with those of high-temperature hydrothermal fluids in modern oceans (pHin-situ = approximately 5) (Seyfried et al. 1991).
Fig. 4

Changes in a dissolved CO2 and b H2 concentrations in the sampled fluids as a function of reaction time

In both experiments, the H2 concentration in the fluid suddenly increased and temporarily reached a high value just after the start of the experiments, then slightly decreased in each experiment (Fig. 4b). Subsequently, the H2 concentration remained almost constant in Exp-350, whereas it gradually increased with time in Exp-250. Finally, it converged to reach a steady state at up to 0.024 mmol/kg in Exp-250 and 2.9 mmol/kg in Exp-350 (Table 2). This trend indicates that the H2 concentration was unstable due to extensive water/rock reactions just after the start of the experiments, and then gradually reached a steady state. The steady-state H2 concentration in Exp-350 was approximately 100 times higher than that in Exp-250, thus indicating that differences in temperature significantly affect hydrogen generation during water-rock reactions.

Alteration products

The XRD analysis revealed that the synthetic komatiites were extensively altered throughout the experiments (Fig. 5). The results show that the alteration product in Exp-250 contains dolomite, serpentine, and a smectite/chlorite mixture (Fig. 5b), whereas calcite, serpentine, and a smectite/chlorite mixture were present as major alteration minerals in Exp-350 (Fig. 5c). The existence of serpentine and carbonate minerals in the run products indicated komatiite was altered by both serpentinization and carbonation. The precipitated carbonate mineral differed for the two temperatures.
Fig. 5

Representative XRD patterns (CuKα) of a the unaltered synthetic komatiite and the alteration products of b Exp-250 and c Exp-350. Dol dolomite, Cal calcite, Sme smectite, Chl chlorite, Ol olivine, and Serp serpentine

Secondary electron imaging (SEI) and EPMA spot analysis provided more detailed information on mineral assemblage and composition (Fig. 6 and Table 3). In Exp-250, EPMA analyses of the products revealed that dolomite crystals contained a relatively high content of iron (3–9 wt.% FeOtotal), whereas the serpentine and smectite/chlorite mixture had FeOtotal contents of 13–15.5 wt.% (Fig. 6a–d; Table 3). Siderite (including approximately 40 wt.% FeOtotal) and olivine were present as minor minerals in the alteration products in Exp-250 (Fig. 6a, c; Table 3). In Exp-350, the EPMA analyses of the alteration products revealed that calcite crystals had a relatively low FeOtotal content (0.1–0.8 wt.%) and that the serpentine and smectite/chlorite mixture had FeOtotal contents of 14–17 wt.% (Fig. 6e; Table 3). Olivine and magnetite were also identified as minor minerals; however, according to the SEI observations in many places over the thin sections, their abundances were likely greater in Exp-350 than in Exp-250. Comparing the compositions of serpentine and the smectite/chlorite mixture between Exp-250 and Exp-350, no large difference in FeOtotal content was identified. On the other hand, on average, the FeOtotal content in Fe-bearing dolomite in Exp-250 was ten times higher than that in calcite in Exp-350 (Table 3).
Fig. 6

Representative SEI images of alteration products (ad) in the experiment at 250 °C and (e, f) in the experiment at 350 °C. Sd siderite, Dol dolomite, Cal calcite, Sme smectite, Chl chlorite, Ol olivine, Serp serpentine, and Mag magnetite

Table 3

Composition of run products (wt.%)

Experiment-250 °C

Experiment-350 °C

Sample

Serp or Sme/Chl mixtureb

Dolomite

Siderite

Serp or Sme/Chl mixtureb

Calcite

Olivinec

Magnetite

SiO2

41.37

40.89

33.38

0.00

0.00

0.00

3.80

38.55

35.52

38.87

0.00

0.00

0.00

41.60

2.76

TiO

0.44

0.43

0.64

0.01

0.05

0.15

2.96

0.52

0.47

0.17

0.00

0.00

0.00

0.01

0.18

Al2O3

7.00

6.88

11.56

0.06

0.00

0.01

0.13

6.82

6.84

6.38

0.07

0.00

0.01

0.06

0.02

Cr2O3

0.47

0.43

0.29

0.11

0.06

0.12

1.46

0.37

0.47

0.39

0.04

0.00

0.02

0.14

0.23

FeOtotal a

13.32

13.65

15.55

3.22

6.66

9.05

39.55

16.84

16.51

13.77

0.18

0.37

0.77

5.33

85.06

MnO

0.16

0.14

0.18

0.53

0.49

0.49

0.12

0.19

0.34

0.20

0.13

0.36

0.72

0.05

0.40

MgO

24.05

23.48

24.10

17.69

17.39

17.43

1.32

23.13

25.21

25.98

0.17

0.74

0.79

51.46

0.36

CaO

0.83

0.81

0.17

26.01

29.47

23.75

0.26

0.28

0.13

0.46

56.29

54.95

58.68

0.32

0.29

Na2O

0.06

0.08

0.03

0.03

0.11

0.04

0.09

0.16

0.07

0.15

0.01

0.00

0.00

0.00

0.02

K2O

0.09

0.11

0.02

0.03

0.02

0.02

0.05

0.04

0.01

0.03

0.01

0.01

0.01

0.02

0.03

NiO

0.17

0.20

0.12

0.00

0.00

0.08

0.10

0.10

0.00

0.16

0.01

0.00

0.05

0.45

0.10

Total

87.96

87.10

86.05

47.67

54.25

51.13

49.84

86.98

85.59

86.56

56.91

56.43

61.04

99.43

89.45

Mg# d

76.30

75.40

73.43

90.73

82.31

77.43

5.63

71.00

73.13

77.08

61.52

78.23

64.50

94.51

0.74

aTotal iron as FeO

bSerpentine or smectite/chlorite mixture

cAlterlation or relict olivine

dMg# = [Mg / (Mg+Fe)] × 100

Dissolved CO2 concentration and pH

The ΣCO2 concentrations in fluids decreased from approximately 400 to 33 mmol/kg in Exp-250 and 171 mmol/kg in Exp-350 (Fig. 4a and Table 2). These differing steady-state ΣCO2 concentrations indicate that carbonate minerals are destabilized with increasing temperature in the komatiite-CO2-H2O system. This trend is consistent with previous basalt experiments under CO2-rich conditions at 250 and 350 °C, where the final ΣCO2 concentrations were 1 mmol/kg at 250 °C and 108 mmol/kg at 350 °C (Shibuya et al. 2013b). Comparing results for the same temperature, the ΣCO2 concentration in our komatiite experiments is higher than that in the basalt experiments of Shibuya et al. (2013b), thus indicating that komatiite has a lesser ability than basalt to absorb CO2 through water-rock reactions. This would be derived from the lower CaO content of komatiite (compared with basalt) because CaO is the main component in the alteration carbonate (mainly calcite or dolomite) in the basalt and komatiite experiments (Shibuya et al. 2013b and the present study).

The final pHin-situ values of fluids in the present experiments (pHin-situ = 4.8 at 250 °C and 5.7 at 350 °C) are clearly lower than those of the previous basalt-CO2-NaCl fluid experiments (pHin-situ = 6.6 at 250 °C and 7.2 at 350 °C) (Shibuya et al. 2013b). This discrepancy is probably derived from the lower pH25 °C of the initial solution (pH25 °C = 4.9) in our experiments than the value of 6.5 in the basalt experiments and/or compositional differences between komatiite and basalt. In general, under CO2-rich and near-neutral pH conditions at room temperature, substantial HCO3 is dissolved in the solution. With increasing temperature, the dissolved HCO3 is converted to CO2(aq) and OH, which elevates the pHin-situ value of the hydrothermal fluid (Shibuya et al. 2010). Although reactions with rocks generally cause fluid pHin-situ to converge near 5.5 (neutral pH at 250–350 °C), previous experiments have revealed that the alkalinization effect due to the increase in temperature under CO2-rich conditions potentially elevates the pHin-situ value beyond a neutral pH value at high temperature, even if the fluid reacts with rocks (Shibuya et al. 2013b). In this study, however, the assumed pH25 °C of the initial solution was 4.9, which was likely too low to form substantial HCO3 in the initial solution at room temperature. This implies that if seawater pH was lower than approximately 5 (CO2aq was the main species), metal-rich black smoker-type hydrothermal fluids would have been generated in the Hadean komatiite hydrothermal systems in contrast to the metal-poor, alkaline hydrothermal fluids in the Archean basalt-hosted system. Although the difference in whole rock composition between komatiite and basalt should be experimentally evaluated, seawater pH was likely an important factor controlling the composition of hydrothermal fluids in the early Earth.

Serpentinization of komatiites

In this section, we compare our experimental results with conditions in natural environments and thermodynamic calculations in order to characterize the reaction occurring in the reaction cell. The EPMA analysis of alteration products revealed the presence of a small quantity of olivine in both experiments, but the amount of olivine in Exp-350 was greater than that in Exp-250. Thermodynamic calculations for serpentinization of olivine and harzburgite (80 wt.% of olivine, 15 wt.% of orthopyroxene, and 5 wt.% of clinopyroxene) indicate that olivine becomes stable and is one of the major minerals at temperatures above approximately 315–350 °C in the peridotite- or olivine-H2O system (e.g., McCollom and Bach 2009; Klein et al. 2013). In analogy with the calculation, the experiments in this study suggest that olivine stabilizes with increasing temperature, even during komatiite serpentinization at 250–350 °C. Nevertheless, olivine was observed as a minor phase even in Exp-350. This is probably due to the composition of komatiite that contains a certain level of Al2O3 would stabilize serpentine and/or chlorite/smectite because smectite becomes stable as bulk Al2O3 content increases in ultramafic rocks such as komatiite (Shibuya et al. 2015). However, it was difficult to determine whether the igneous olivine still remains in the absence of any reaction with a fluid or if the alteration olivine was newly created through serpentinization.

Brucite commonly occurs in serpentinized rocks in modern oceanic plates (e.g., Bach et al. 2006; Klein et al. 2014) but was not found in the alteration products in the present experiments. The absence of brucite in Exp-350 is consistent with the thermodynamic calculations for the serpentinization of olivine and harzburgite under CO2-poor conditions, since brucite is destabilized at temperatures above ~350 °C (McCollom and Bach 2009; Klein et al. 2013). On the other hand, the instability of brucite in Exp-250 is likely due to the CO2-rich condition in the experiment because previous experiments revealed that Mg-bearing carbonate mineral is more stable than brucite under CO2-rich conditions (Zhao et al. 2010; Hövelmann et al. 2012; Klein and McCollom 2013). Alternatively, high dissolved Si content in the fluid (3.7–4.1 mmol/kg) possibly prohibited brucite formation, which is supported by thermodynamic calculations (e.g., Klein and McCollom 2013).

Carbonation of komatiites

The steady-state ΣCO2 concentrations in both experiments indicate that the amount of carbonate mineral was greater in Exp-250 than in Exp-350 (Fig. 4a and Table 2). The formation of dolomite in Exp-250 is characteristically different from the basalt-CO2 experiments that yielded only calcite as carbonate in previous studies at 250 and 350 °C (Shibuya et al. 2013b). Comparing the experimental conditions of this study to those in the previous study, compositional differences between komatiite and basalt and/or the difference in initial pH (4.9 in this study, and 6.5 in Shibuya et al. (2013b)) likely caused the difference in precipitated carbonate species in the alteration products at 250 °C. Furthermore, it was revealed that the precipitated carbonate species strongly affected the Mg concentration in the hydrothermal fluid, i.e., Mg concentration in Exp-250 (carbonate as dolomite) is 30–40 times higher than that in Exp-350 (carbonate as calcite), while fluid Ca concentrations in both experiments are similar.

Previously, carbonate species were reported to correlate with the concentration ratio between Mg and Ca (mMg2+/(mMg2+ + mCa2+)) and the temperature of the fluid (Fig. 7) (Rosenberg and Holland 1964; Rosenberg et al. 1967; Tribble et al. 1995). In Exp-350, carbonate was precipitated as calcite, whereas the mMg2+/(mMg2+ + mCa2+) value of the fluid was approximately 0.03; thus, it clearly falls within the stability field of calcite. On the other hand, the mMg2+/(mMg2+ + mCa2+) value of the fluid in Exp-250, where dolomite precipitated, was approximately 0.43, which is near the dolomite/magnesite stability boundary. Therefore, the extremely high fluid Mg concentration in Exp-250 is consistent with the precipitation of dolomite in the alteration product. However, it is difficult to discuss the anteroposterior relation between dolomite formation and enrichment of Mg in fluids because the decrease in ΣCO2 concentration and the Mg enrichment in the fluid were already confirmed at the first sampling. It is widely accepted that seafloor hydrothermal systems play a significant role as Mg sinks in the modern ocean because basaltic oceanic crusts incorporate seawater Mg into alteration minerals such as smectite and chlorite through hydrothermal alteration at near mid-ocean ridges (e.g., Alt 1995). In addition, Charlou et al. (2002) reported that ultramafic rock-hosted high-temperature hydrothermal systems provide fluids with little Mg into seawater and generally act as Mg sinks in the ocean. On the other hand, weathering of ultramafic rocks exposed at seafloor below 150 °C and hydrothermal circulation at high water/rock ratios provide Mg into seawater (Snow and Dick 1995). Our results hypothesize that dolomite formation during the serpentinization of komatiite would be a source of Mg for the Hadean ocean, even when the temperature range of hydrothermal reaction zones was near 250 °C. Further experiments will be necessary to justify this hypothesis.
Fig. 7

The mMg2+/(mMg2+ + mCa2+) values of hydrothermal fluids in Exp-250 and Exp-350 plotted on the calcite-dolomite-magnesite stability fields estimated from experiments at 275–420 °C and their extrapolation to 0 °C (modified after Tribble et al. 1995). The dotted and solid curves are for the solutions with 2 and 1 M of Ca-Mg chloride, respectively. The blue circle and pink diamond indicate the values in Exp-250 and Exp-350, respectively

The dolomite crystals formed in this study contain up to 9 wt.% FeOtotal in Exp-250 because Fe substitutes Mg in dolomite during its precipitation owing to the similarity in ion radius between Fe and Mg (Mg2+: 0.72 Å; Fe2+: 0.61 Å) (Jia 1991). A similar Fe-bearing carbonate mineral was also reported in a hydrothermal alteration experiment for komatiite conducted under CO2-rich conditions (water 50.1 mg + CO2 7.3 mg: about 3300 mmol/kg at the starting condition) at 300 °C and 500 bars (Hao and Li 2015). Furthermore, Klein and McCollom (2013) reported the precipitation of Fe-bearing magnesite (FeO = 6.78 wt.%) through the serpentinization of olivine under CO2-rich conditions at 230 °C. Although differences in the composition of the starting materials between the previous studies and our experiments would have caused the difference in carbonate species, it seems likely that carbonate minerals formed during serpentinization under CO2-rich conditions contain substantial FeO, probably at temperatures below approximately 300–350 °C.

Effect of seawater CO2 on H2 concentration in fluid

The maximum steady-state fluid H2 concentrations were 0.024 mmol/kg in Exp-250 and 2.9 mmol/kg in Exp-350 (Fig. 4b), and the results are consistent with the analysis of solid alteration products. Magnetite, whose formation is related to H2 generation (McCollom and Bach 2009), was observed in Exp-350 but not in Exp-250. Nevertheless, the H2 concentrations in both experiments were clearly lower than the steady-state H2 concentration (~20 mmol/kg) in fluid generated by the reactions between komatiite and a CO2-free NaCl fluid at 300 °C and 500 bars (Shibuya et al. 2015). Although direct comparison is difficult owing to the difference in temperature conditions, two possible factors explain the differences between the previous and present experiments in terms of the steady-state H2 concentrations from the hydrothermal reaction of komatiite.

The first factor is the temperature dependence of the stability of magnetite in ultramafic systems. At temperatures below 315–350 °C, the amount of magnetite formation increases with increasing temperature, but it drastically decreases at temperatures above 315–350 °C because olivine becomes stable (McCollom and Bach 2009; Klein et al. 2013). Moreover, Klein et al. (2013) reported that within the temperature range of 25–400 °C, the amount of hydrogen generation caused by the reaction with olivine (Fo 90) reaches a maximum at 322 °C. Thus, the temperature dependence of the stability of olivine and magnetite would be a primary cause of the difference in fluid H2 concentrations between the previous CO2-free komatiite experiment at 300 °C and our experiments.

The other factor is the presence of CO2 in the experimental system. Based on thermodynamic calculations simulating the relative amount of magnetite formation at 50–400 °C under 500 bars, Klein et al. (2013) expected that the amount of hydrogen generation caused by the reaction with olivine under CO2-free conditions is greater at 250 °C than at 350 °C. In our experiments, however, the steady-state fluid H2 concentration was higher in Exp-350 than in Exp-250. In previous experiments that simulated the reactions between olivine and a CO2-rich fluid at 230 °C and 35 MPa, ferrous iron-bearing talc and magnesite were generated, while magnetite formation and H2 generation were limited (Klein and McCollom 2013). Under CO2-free conditions, the iron is only partitioned into magnetite and ferrous iron-bearing silicate in alteration products,
$$ \mathrm{a}{\left(\mathrm{F}\mathrm{e}\mathrm{O}\right)}_{\mathrm{rock}}+{\mathrm{bH}}_2\mathrm{O}\to \mathrm{d}{\left({\mathrm{Fe}}_2{\mathrm{O}}_3\right)}_{\mathrm{rock}}+d{\mathrm{H}}_2+\left(\mathrm{a}\hbox{-} 2d\right){\left(\mathrm{F}\mathrm{e}\mathrm{O}\right)}_{\mathrm{rock}.}. $$
In contrast, the interaction between CO2-rich fluid and rock leads to carbonate mineral precipitation and ferrous iron incorporation therein. The reaction occurring under a CO2-rich condition is the following:
$$ \mathrm{a}{\left(\mathrm{F}\mathrm{e}\mathrm{O}\right)}_{\mathrm{rock}}+{\mathrm{bH}}_2\mathrm{O}+{\mathrm{cCO}}_2\to \mathrm{d}{\left({\mathrm{Fe}}_2{\mathrm{O}}_3\right)}_{\mathrm{rock}}+{\mathrm{dH}}_2+\mathrm{e}{\left(\mathrm{F}\mathrm{e}\mathrm{O}\right)}_{\mathrm{rock}}+\left(\mathrm{a}\hbox{-} 2d\hbox{-} \mathrm{e}\right){\left({\mathrm{Fe}\mathrm{CO}}_3\right)}_{\mathrm{rock}}+\left(\mathrm{c}\hbox{-} \mathrm{a}+2d+\mathrm{e}\right){\mathrm{CO}}_2. $$

In both equations, the value of “d” is the amount of H2 generation and depends on temperature. Therefore, it was suggested that the lack of H2 generation is due to the direct incorporation of ferrous iron into carbonate minerals without its oxidation during the hydrothermal alteration. In our experiments, although a large difference in the composition of serpentine and smectite/chlorite mixture was not identified between Exp-250 and Exp-350, the Fe content of dolomite in Exp-250 (3–9 wt.% FeOtotal) was clearly higher than that of calcite in Exp-350 (Table 3). The amount of generated magnetite in Exp-350 was greater than that in Exp-250. These data suggest that the amount of ferrous iron incorporated into the carbonate mineral strongly affected the extent of magnetite formation and resulting hydrogen generation in the fluid. Although such a trend in the effect of CO2 on H2 generation has been reported in some previous experiments at temperatures below 230 °C (Jones et al. 2010; Klein and McCollom 2013; Neubeck et al. 2014), our experiments show that the suppression of H2 generation by the presence of CO2 in the system occurs even at higher temperatures (250 °C) and in the CO2-H2O-komatiite system. In addition, a small amount of ferrous iron incorporated into calcite (up to 0.8 wt.%) may have also slightly suppressed magnetite formation in Exp-350.

Implications for the Hadean H2-rich hydrothermal environments

It has been considered that the serpentinization of komatiite would have served as the most ubiquitous geological process for hosting H2-rich hydrothermal environments in the Hadean ocean (e.g., Takai et al. 2006). The present experiments constrained the potential for H2 generation during the serpentinization of komatiite under CO2-rich conditions. This work provides further important insights into our interpretation of hatcheries for the emergence and early evolution of life in the Hadean Earth. Exp-350 showed that the fluid H2 concentration increased up to 2.9 mmol/kg during the serpentinization and carbonation of komatiite. This value is lower than the maximum H2 concentration of the vent fluids in modern peridotite-hosted hydrothermal systems; the H2 and ΣCO2 concentrations in modern peridotite-hosted hydrothermal systems are 12 and 10.1 mM at Logatchev field, 16 and 16 mM at Rainbow field (Charlou et al. 2002), and <1–15 mmol/kg and <0.8 mmol/kg at Lost City field (Kelley et al. 2005), respectively. However, the H2 concentration in Exp-350 broadly falls within the range of the fluid H2 concentrations in the Kairei field hosted by both basalt and troctolite (H2 = 2.5–8.5 mmol/kg and ΣCO2 < 10.1 mmol/kg; Takai et al. 2004; Gallant and Von Damm 2006; Kumagai et al. 2008; Nakamura et al. 2009). This suggests that in the Hadean, the komatiite still had great potential to generate H2-rich hydrothermal fluids at high-temperature, even under CO2-rich conditions. Furthermore, considering that hydrogenotrophic methanogens—long believed to be one of the most probable candidates for the earliest living forms and primary producers on Earth (e.g., Takai et al. 2006; Martin et al. 2008)—indeed dominate in the microbial communities associated with hydrothermal fluids in the Kairei field (Takai et al. 2004), the komatiite-hosted seafloor hydrothermal systems would have fully prepared the energetic basis of hatcheries for the emergence and early evolution of hydrogenotrophic living forms in the Hadean ocean.

In contrast, the fluid H2 concentration in Exp-250 (0.024 mmol/kg) was much lower than that in Exp-350 and is comparable to the H2 concentration level in the typical fluids of modern basalt-hosted hydrothermal fields (Charlou et al. 1996, 2000; Gallant and Von Damm 2006). Given that the H2 generation potential of ultramafic rocks decreases with decreasing temperature within the unstable region of olivine (McCollom and Bach 2009; Klein et al. 2013), the serpentinization of komatiite under CO2-rich conditions might not have afforded the H2-rich hydrothermal fluids at temperatures lower than 250 °C. More importantly, carbonate minerals become stable with decreasing temperature in H2O-CO2-rock systems (e.g., Shibuya et al. 2013b), which indicates that the carbonate formation and resulting limitation of H2 generation would occur more significantly at lower temperatures (e.g., <100 °C). Therefore, the possible H2-rich hydrothermal environments may have been created only in the deep ocean floor because high hydrostatic pressure can elevate the temperature of hydrothermal reaction zones (fluid-rock reaction zones). In other words, other potential candidates, such as the hydrothermal systems in on-land and shallow submarine environments, may not have fully prepared the energetic basis (abundant H2 availability) of hatcheries for the emergence and early evolution of life in the Hadean Earth due to low-temperature hydrothermal reactions.

Conclusions

High-temperature and high-pressure experiments using komatiite and a CO2-rich NaCl solution revealed that different hydrothermal reaction temperatures caused differences in the carbonate species in alteration products (iron-rich dolomite at 250 °C and calcite at 350 °C) during the serpentinization of komatiites. The hydrothermal fluid coexisting with dolomite at 250 °C showed high Mg concentrations (up to ca. 40 mmol/kg), which is markedly higher than those in typical modern high-temperature hydrothermal fluids (generally less than 1 mmol/kg). This suggests that, in contrast to modern seafloor hydrothermal systems, the dolomite-bearing komatiite-hosted hydrothermal systems may have served as a source of Mg in the Hadean ocean. The steady-state H2 concentration in Exp-350 (up to 2.9 mmol/kg) was approximately 100 times higher than that in Exp-250 (up to 0.024 mmol/kg) because the amount of ferrous iron incorporated into carbonate was greater in Exp-250 than in Exp-350. These results perhaps suggest that only high-temperature komatiite-hosted hydrothermal systems have the potential to generate H2-rich hydrothermal environments, even in the CO2-abundant Hadean ocean. Further experiments based on a more precise estimation of the ancient atmospheric/oceanic CO2 levels from geological records will clarify the geochemical nature of komatiite-hosted hydrothermal environments and other potential hatcheries for the emergence and early evolution of life in the Hadean Earth because the ΣCO2 concentration in the fluid may significantly affect the mineral assemblage of serpentinized komatiites and hydrothermal reactions taking place under various conditions.

Abbreviations

BSE: 

Backscattered electron

EPMA: 

Electron probe microanalyzer

Exp-250: 

Experiment at 250 °C

Exp-350: 

Experiment at 350 °C

QFM: 

Quartz-fayalite-magnetite

SEI: 

Secondary electron imaging

XRD: 

X-ray diffraction

XRF: 

X-ray fluorescence

Declarations

Acknowledgements

We thank two anonymous reviewers for many helpful comments that greatly improved our manuscript. We appreciate Prof. Hodaka Kawahata for editorial handling. We are grateful to Y. Ueno for discussions.

Funding

This work was supported by Grants-in-Aid for Scientific Research from the Japan Society for Promotion of Science (JSPS) (No. 15K13583) and the Ministry of Education, Culture, Sports, Science and Technology (MEXT), Japan.

Authors’ contributions

HU, TS, KT, and SM proposed and designed the study. HU and TS carried out the experiments. HU, TS, and MS prepared the starting materials. HU, TS, MS, and YS analyzed the fluid samples and solid materials. All authors interpreted the data, read, and approved the final manuscript.

Competing interests

The authors declare that they have no competing interests.

Open AccessThis article is distributed under the terms of the Creative Commons Attribution 4.0 International License (http://creativecommons.org/licenses/by/4.0/), which permits unrestricted use, distribution, and reproduction in any medium, provided you give appropriate credit to the original author(s) and the source, provide a link to the Creative Commons license, and indicate if changes were made.

Authors’ Affiliations

(1)
Department of Earth and Planetary Sciences, Tokyo Institute of Technology
(2)
Laboratory of Ocean-Earth Life Evolution Research (OELE), Japan Agency for Marine-Earth Science and Technology (JAMSTEC)
(3)
Research and Development Center for Submarine Resources, Japan Agency for Marine-Earth Science and Technology (JAMSTEC)
(4)
Project Team for Development of New-generation Research Protocol for Submarine Resources, Japan Agency for Marine-Earth Science and Technology (JAMSTEC)
(5)
Department of Subsurface Geobiological Analysis and Research (SUGAR), Japan Agency for Marine-Earth Science and Technology (JAMSTEC)
(6)
Earth-Life Science Institute (ELSI), Tokyo Institute of Technology

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