The subseafloor thermal gradient at Iheya North Knoll, Okinawa Trough, based on oxygen and hydrogen isotope ratios of clay minerals

Journal of Volcanology and Geothermal Research 384 (2019) 263–274

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The subseafloor thermal gradient at Iheya North Knoll, Okinawa Trough, based on oxygen and hydrogen isotope ratios of clay minerals Youko Miyoshi a,b,⁎, Jun-ichiro Ishibashi a, Seiichiro Uehara a, Kazuhiko Shimada a, Kevin Faure c a b c

Department of Earth and Planetary Sciences, Graduate School of Science, Kyushu University, Japan Research Institute for Geo-Resources and Environment, National Institute of Advanced Industrial Science and Technology (AIST), Japan Department of Earth Structure and Processes, GNS Science, New Zealand

a r t i c l e

i n f o

Article history: Received 2 April 2019 Received in revised form 19 July 2019 Accepted 19 July 2019 Available online 23 July 2019 Keywords: Clay mineral Seafloor hydrothermal system Hydrothermal alteration Oxygen isotope Hydrogen isotope Formation temperature

a b s t r a c t We measured the oxygen and hydrogen isotope ratios of smectite, chlorite, and illite in hydrothermally active sediments at the Iheya North Knoll, Okinawa Trough, obtained by seafloor drilling of active hydrothermal fields during the Integrated Ocean Drilling Program (IODP) Expedition 331. Formation temperatures calculated from oxygen isotope data were compared to the downhole temperatures that were measured during the same expedition. The δ18O values ranged from +9.9‰ to +13.3‰ for smectite, +1.3 to +3.0‰ for chlorite or chlorite/smectite mixed-layer mineral, and +2.1‰ to +5.1‰ for illite or illite/smectite mixed-layer mineral with minor chlorite. For samples of illite that contained minor chlorite, δ18O values for illite or illite/smectite mixed-layer mineral of +4.4‰ to +6.1‰ were calculated, based on the proportion of the two clay minerals calculated from chemical composition of the individual mineral and the mixture determined by TEM-EDS and EPMA. We estimated the formation temperatures of the minerals using oxygen isotope fractionation data and assumed values of 0 to +1.2‰ for the fluid that was equilibrated with the minerals, based on the reported values of the seawater and vent fluid measured in previous studies. The temperature range calculated was 111–175 °C for a shallow layer less than about 25 m below seafloor (mbsf) that contained smectite and a minor kaolinite component, 177–297 °C for a deeper layer between about 25 and 40 mbsf that was dominated by chlorite or chlorite/ smectite mixed-layer mineral, and 220–314 °C for the deepest layer between about 40 and 120 mbsf that contained illite or illite/smectite mixed-layer mineral with less chlorite. The temperatures of mineral formation calculated from the oxygen stable isotope data were slightly higher at some depths, but both calculated and measured temperatures demonstrated a steep, step-wise thermal gradient below the active hydrothermal field, where temperatures of N200 °C were determined at as little as 50 mbsf. Hydrogen isotope values mostly ranged between −40 and −55‰ and did not vary significantly or systematically between the layers, except in the clay fractions containing smectite above 25 mbsf that had relatively low values between −65 and −72‰. Hydrogen isotope results of the clay minerals were problematic to interpret and not consistent with interpretation of the oxygen data. Calculated temperatures from illite and smectite were mostly negative, whereas temperatures calculated from chlorite were more positive (average 145 °C), but lower than those from the oxygen isotope data. Additional work is still required to determine the hydrogen isotope fractionation between clay minerals and water with respect to temperature. © 2019 Elsevier B.V. All rights reserved.

1. Introduction Stable isotope values of clay minerals provide information about the temperature of formation, based on the temperature dependency of the isotope fractionation factor. Oxygen isotope ratios of clay minerals reflect the temperature and oxygen isotope ratios of fluids, because clay minerals generally form in isotopic equilibrium with their surrounding ⁎ Corresponding author at: Research Institute for Geo-Resources and Environment, National Institute of Advanced Industrial Science and Technology (AIST), Central-7, 1-11 Higashi, Tsukuba, Ibaraki 305-8567, Japan. E-mail address: [email protected] (Y. Miyoshi).

https://doi.org/10.1016/j.jvolgeores.2019.07.017 0377-0273/© 2019 Elsevier B.V. All rights reserved.

environment (Savin and Hsieh, 1998). It is well established that the oxygen isotope ratios of clay minerals in active geothermal fields generally decrease with increasing depth and temperature, assuming that the oxygen isotope ratio of the circulating fluid does not change significantly (Inoue et al., 2004; Libbey et al., 2013). Estimating formation temperature based on oxygen isotope ratios is feasible for clay minerals collected from seafloor hydrothermal fields, especially for those in sediment-rich environments where abundant clay minerals are usually formed (e.g., Zierenberg and Shanks III, 1994; Buatier et al., 1995; Marumo and Hattori, 1999; Lackschewitz et al., 2000; Lackschewitz et al., 2004; Dekov et al., 2008; Marumo et al., 2008; Miyoshi et al., 2013). These studies have demonstrated that

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temperatures were high enough for the formation of certain clay minerals. There still is a need for a capability to determine clay mineral formation temperature in seafloor hydrothermal fields, because conducting downhole temperature measurements are not as common as in exploration of on-land geothermal systems. Lackschewitz et al. (2000) studied clay minerals collected from hydrothermal deposits in the Middle Valley, Juan de Fuca Ridge, during Ocean Drilling Program (ODP) Leg 169 and showed that the estimated formation temperature of chlorite and chlorite/smectite mixed-layer minerals based on their oxygen isotope ratios were similar to the downhole temperature measured during the same leg at a neighboring borehole located ~50 m away. We analyzed oxygen and hydrogen isotope ratios of clay minerals in sediment that were collected from an active hydrothermal field at the Iheya North Knoll in the Okinawa Trough back-arc basin during Integrated Ocean Drilling Program (IODP) Expedition 331, conducted in September 2010. With this data we assess the subseafloor thermal profile of the hydrothermal fields. During the same expedition, direct temperature measurements were attempted and in situ temperature information was obtained at seven different depths (Takai et al., 2011). This dataset enables us to compare the estimated formation temperatures of clay minerals with direct downhole temperature measurements. 2. Background 2.1. Geological setting of the drill site The Okinawa Trough is a back-arc basin located between the Ryukyu island-arc and the Asian continent (Fig. 1a). The seafloor in the Okinawa Trough is covered with sediment that has a thickness of approximately 2 km (Tsuji et al., 2012; Letouzey and Kimura, 1986). The study area was located at the Iheya North Knoll which is one of abundant small knolls in the middle Okinawa Trough. At the east side of the Iheya North Knoll, nine hydrothermal vents with sulfide mounds have been discovered (Kawagucci et al., 2011). The North Big Chimney (NBC) mound, N30 m in height and associated with the vigorous venting of clear fluid with a maximum temperature of 311 °C, appears to mark the center of hydrothermal activity at the Iheya North Knoll (Nakagawa et al., 2005). In September 2010, IODP Expedition 331 was conducted in the Iheya North Knoll using the deep-sea drilling vessel Chikyu. The expedition involved drilling at five sites on the east slope of the Iheya North Knoll, with three sites (C0013, C0014, and C0016) being close to the NBC mound (Fig. 1b). Eight holes (C0013A-C0013H) were drilled at Site C0013, which is located about 100 m east of the NBC mound (Fig. 1b). Cores were recovered from all holes except hole C0013A. Hole C0013E was the deepest (54.5 m below the seafloor [mbsf]) and was cased down to 40.2 mbsf and fixed with a corrosion cap (open outlet pipe) mounted on the guide base (Takai et al., 2011; Kawagucci et al., 2013; Nozaki et al., 2016). The core liner melted at 12.5 mbsf at hole C0013C, which indicated that the in situ temperature was at least 82 °C even at this depth (Takai et al., 2011). Discharging of hydrothermal fluid was observed at hole C0013E immediately after drilling started, for which the temperature was confirmed to be N250 °C based on observations of thermoseal temperature-sensitive strips taped on the corrosion cap outlet pipe (Takai et al., 2011). Seven boreholes (C0014A-C0014G) were drilled at site C0014, which is located about 450 m east of the NBC mound (Fig. 1b). Borehole C0014G was the deepest (136.7 mbsf) and was cased down to 117.8 mbsf and fixed with a corrosion cap (Takai et al., 2011; Kawagucci et al., 2013; Nozaki et al., 2016). Hydrothermal fluid was discharged from the holes after penetrating hard layers at depths of 35–44.4 mbsf in hole C0014B, 25.5–35 mbsf in hole C0014E, and 37.7–47.2 mbsf and 89.2–93.7 mbsf in borehole C0014G (Takai et al., 2011). The temperatures of these discharging fluids were also

Fig. 1. (a) Location of the Iheya North Knoll field, Okinawa Trough (after Suzuki et al., 2008). (b) Detailed map of the Iheya North Knoll field with the locations of sites C0016, C0013, and C0014 (after Takai et al., 2011). (c) Cross-section of the seafloor of the Iheya North Knoll with the locations of the drilling sites (after Miyoshi et al., 2015). The layers were determined on the basis of clay mineral assemblages (Miyoshi et al., 2015) (see text for details).

determined to be N240 °C based on observations of thermoseal strips, as in the case of the holes in site C0013 (Takai et al., 2011). Two boreholes (holes C0016A and C0016B) were drilled at site C0016, which is located near the NBC mound (Fig. 1b), but only one core was recovered (C0016B). Borehole C0016A was drilled directly into the top of the NBC mound. Borehole C0016B was drilled about 20 m west of the NBC mound and borehole C0016A. Borehole C0016B penetrated to 45 mbsf, but the total recovery of cores was only 2.095 m (Takai et al., 2011). Black smoke discharged vigorously from borehole C0016B when the layer at 27–45 mbsf was penetrated (Takai et al., 2011) and the temperature of the discharging fluid was 275 °C five months after the expedition (Kawagucci et al., 2013).

Y. Miyoshi et al. / Journal of Volcanology and Geothermal Research 384 (2019) 263–274

2.2. Summary of clay mineralogy Miyoshi et al. (2015) documented the clay mineralogy in detail for three sites (C0013, C0014, and C0016) close to the NBC mound. Sediments were classified into Layers 0–4 with increasing depth, on the basis of the assemblages of clay minerals. Layer 0 consisted mainly of quartz and volcanic glass and lacked clay minerals. Layer 1 contained smectite, kaolinite, and illite/smectite mixed-layer mineral. Layer 2 contained chlorite, chlorite/smectite mixed-layer mineral, and corrensite (regular 1:1 mixed-layer mineral of chlorite and smectite). Layer 3 contained chlorite (or chlorite/smectite mixed-layer mineral) and illite (or illite/smectite mixed-layer mineral). From here on the illite/smectite mixed-layer mineral is

265

abbreviated as illite/smectite, and the chlorite/smectite mixedlayer mineral as chlorite/smectite. The depth of the clay layers differed between sites C0013 and C0014 (Fig. 1c). Layer 0 was not found at site C0013. The boundary between Layers 0 and 1 was ~12 mbsf at site C0014. The boundary between Layers 1 and 2 was ~6.5 mbsf at site C0013 and ~25 mbsf at site C0014. The boundary between Layers 2 and 3 was ~28 mbsf at site C0013 and ~40 mbsf at site C0014. The boundaries were unclear at site C0016 because of poor core recovery, but several large pieces of core were recovered. Corrensite and illite were found at shallow depth (b1 mbsf), corrensite at moderate depth (9.1 mbsf), and chlorite and illite at the maximum depth (27 mbsf).

Table 1 Oxygen and hydrogen isotope ratios of clay fraction samples. (a) Site C0013 Sample ID Hole C0013B C0013B 1T-1 80-82 Hole C0013C C0013C 1H-3 22-24 C0013C 1H-14 49-51 Hole C0013D C0013D 1H-1 57-67 C0013D 1H-2 68-80 C0013D 1H-3 107.5-117.5 C0013D 2H-2 16-18 C0013D 2H-6 14-16 Hole C0013E C0013E 6X-CC 22-24 C0013E 7L-2 87.5-99.5 C0013E 9X-CC 8-10

Depth (mbsf)

Layera

Dominant clay minerala

δ18O (‰) 9.6b

δD (‰) −40.6

0.8

1

Kaolinite, illite

4.7 11.1

1 2

Smectite Corrensite

3.6 4.5 5.9 13.7 18.0

1 1 1 2 2

I/Sm Kaolinite, smectite Smectite, chlorite Chlorite Chlorite

6.3 8.0b 8.2 1.6 2.4

−40.9 −49.5 n.a. −48.4 −48.9

23.2 27.2 45.1

2 2 3

Chlorite Chlorite Chlorite, illite

2.4 3.0 3.1

−47.3 −46.3 −43.5

11.4 4.2

−51.9 −43.4

(b) Site C0014 Sample ID Hole C0014B C0014B 2H-7 40-50 C0014B 2H-10 20-30 C0014B 3H-2 52-62 C0014B 3H-5 20-30 C0014B 4H-2 35-45 C0014B 4H-4 35-45 C0014B 4H-7 111-121 C0014B 5H-13 53-65 C0014B 5H-15 80-91 Hole C0014G C0014G 3H-2 50-65 C0014G 3H-5 32-47 C0014G 4H-5 45-60 C0014G 4H-7 0-15 C0014G 4H-10 0-15 C0014G 4H-11 80-95 C0014G 6H-3 47-59 C0014G 9X-2 25-40 C0014G 12H-3 15-25 C0014G 14T-2 15-30 C0014G 16T-1 52-62 C0014G 18T-4 20-35 C0014G 20T-1 43-63 C0014G 23X-1 12-24 C0014G 25T-1 35-45

Depth (mbsf)

Layera

Dominant clay minerala

δ18O (‰)

δD (‰)

12.8 15.1 17.4 20.1 26.4 29.2 33.0 42.0 44.7

1 1 1 1 2 2 2 3 3

Kaolinite, smectite Kaolinite, smectite Smectite Smectite Chl/Sm Chl/Sm Chl/Sm I/Sm, chlorite I/Sm, chlorite

12.5b 8.7b 11.8 9.9 2.9 2.0 2.1 4.7 5.1

−52.2 −52.0 −65.8 −55.3 −47.8 −47.0 −45.1 −52.4 −46.0

19.8 22.9 30.9 32.4 35.5 37.3 48.8 55.9 65.3 71.8 76.4 87.2 94.1 104.8 114.1

1 1 2 2 2 2 3 3 3 3 3 3 3 3 3

Smectite Smectite Chl/Sm Chl/Sm Chl/Sm Chl/Sm I/Sm, chlorite I/Sm, chlorite I/Sm, Chl/Sm I/Sm, Chl/Sm I/Sm, Chl/Sm Illite, chlorite Illite, chlorite Illite, chlorite Illite, chlorite

13.3 11.7 2.9 1.3 2.6 1.5 4.3 4.0 4.0 3.5 3.8 2.1 4.2 4.2 4.2

−71.8 −56.6 −49.0 −40.6 −38.4 −44.8 −49.9 −49.1 −51.3 −49.9 −43.9 −53.2 −47.2 −46.9 −49.4

(c) Site C0016 Sample ID

Depth (mbsf)

Dominant clay minerala

δ18O (‰)

δD (‰)

C0016B 1L-1 37-47 C0016B 3L-1 23-26

0.4 27

Corrensite, illite Chlorite, illite

4.7 3.7

−48.0 −50.4

n.a.: not analyzed. Abbreviations for clay minerals; Chl/Sm: chlorite/smectite mixed-layer mineral, I/Sm: illite/smectite mixed-layer mineral. a Miyoshi et al. (2015). b Yanagawa et al. (2017).

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3. Methods 3.1. Sample preparation This study used the same clay fraction samples as those studied by Miyoshi et al. (2015). The clay fraction samples were obtained from sediment samples by elutriation. Sediments were disaggregated in distilled water with a spoon and ultrasonic bath and rinsed several times to remove dissolved salts. The clay fraction samples (b2 μm) were obtained by suspending the sediments in distilled water for 5 h to allow the larger particles to settle. Clay minerals in the clay fraction samples were identified by X-ray diffraction (XRD), as reported in Miyoshi et al. (2015). We selected clay fraction samples that included one or two clay minerals. 3.2. Oxygen and hydrogen isotope measurements Oxygen and hydrogen isotope ratios of the clay fraction samples were determined at GNS Science. Prior to the isotope measurement, free Fe-oxides were removed from the clay fraction following the method of Mehra and Jackson (1960). Sodium-citrate solution (0.3 mol/L, 10 mL) and NaHCO3 solution (1 mol/L, 1.25 mL) were added to the clay fraction sample (b1 g). The mixture was heated at 80 °C in a water bath, then Na2S2O4 (0.25–0.3 g) was added. The mixture was stirred occasionally for 15 min, then NaCl solution (2.5–3 mL) was added to promote flocculation. The suspension was centrifuged three or four times to wash the sample. The sample was dried and the mineral in the sample was identified again by XRD. For measurement of oxygen isotopes, oxygen was extracted from clay fraction samples (b100 μg) using a CO2 laser and BrF5 following the method of Sharp (1990). Samples and standards were heated to 150 °C overnight prior to loading into the vacuum extraction line. Samples were then held under vacuum at room temperature for about 6 h and blank BrF5 runs were performed until the background yield was b0.2 μmoles oxygen. Oxygen yields were recorded and converted to CO2 gas, which was analyzed using a Geo20-20 mass spectrometer. All oxygen isotope results are reported with respect to Vienna Standard Mean Ocean Water (VSMOW). Oxygen isotope ratios were normalized to the international quartz standard NBS28 using a value of +9.6 per mil (‰), or UWG-2 garnet using a value of +5.8‰. Repeat analyses of standards yielded oxygen values that varied by b0.15‰. For measurement of hydrogen isotopes, samples were pyrolyzed at 1450 °C in silver capsules. Samples were analyzed using a HEKAtech (a) depth (mbsf) 0 0

high temperature elemental analyzer coupled with a GV Instruments IsoPrime mass spectrometer. All samples were analyzed in triplicate. All hydrogen isotope results are reported with respect to VSMOW, normalized to the international standards IAEA-CH-7, NBS30, and NBS22 with reported hydrogen isotope ratios of −100‰, −66‰, and −118‰, respectively. The precision of the standards was ±1.5‰ (1σ). 3.3. Chemical analysis The chemical composition of representative clay fraction samples were determined by the Cu-EPMA method proposed by Miyoshi et al. (2017). The Cu-EPMA method gives the total chemical composition of the clay fraction and can be used in place of X-ray fluorescence (Miyoshi et al., 2017). The clay fraction samples were pasted onto Cu plates and chemical analysis was conducted at five or more points in the clay fraction sample using an electron probe micro-analyzer (EPMA) equipped with a wavelength dispersive spectrometer (WDS). We used an EPMA (JEOL JXA-8530F) at Kyushu University. The chemical compositions were normalized to 100 wt% in total. Detailed analytical conditions were reported in Miyoshi et al. (2017). The chemical composition of individual clay mineral particles was determined with a transmission electron microscope (TEM) equipped with an energy dispersive spectrometer (EDS). We used the TEM-EDS (JEOL JEM-2010FEF; conducted at 200 kV) in the Research Laboratory for High Voltage Electron Microscopy (HVEM), Kyushu University. Samples for the TEM-EDS analysis were settled onto a carbon-coated copper grid after powdered clay fractions were dispersed in alcohol by ultrasonication. Chemical composition was measured using a beam with a diameter of ~0.5 μm, and by referring to the calculated kfactors of the Cliff-Lorimer method. The chemical compositions were normalized to 100 wt% in total. Detailed analytical conditions were reported in Miyoshi et al. (2015). 4. Results and discussion 4.1. Oxygen and hydrogen isotope ratios of clay fractions Results of the oxygen and hydrogen isotope measurements done on the clay fraction are listed in Table 1 and plotted against depth in Figs. 2 and 3. The oxygen isotope values were relatively high in Layer 1 and low in Layer 2 at site C0013 and similar at site C0014 (Fig. 2). In Layer 3 at site C0013 (C0013E 9X-CC 8-10) the oxygen isotope values were similar to those in Layer 2, but the values in Layer 3 at site C0014 were slightly (c)

(b)

5

δ18O (‰, VSMOW) 10 15

20

depth (mbsf) 0 0

18

5

δ O (‰, VSMOW) 10 15

Layer 1

Layer 0

Layer 2

Layer 1

20

depth (mbsf) 0 0

Layer 1

20

40

60

60

smectite kaolinite + illite kaolinite + smectite smectite + chlorite

Layer 2

40 Layer 3

Layer 3

I/Sm corrensite chlorite or Chl/Sm illite or I/Sm + chlorite or Chl/Sm

δ18O (‰, VSMOW) 10 15 Layer 0

Layer 2

40

5

Layer 3

60

80

100

120 Fig. 2. Depth profiles of oxygen isotope ratios of clay fraction samples. (a) Site C0013. (b) Site C0014 (hole B). (c) Site C0014 (hole G).

Y. Miyoshi et al. / Journal of Volcanology and Geothermal Research 384 (2019) 263–274

(b)

(a) depth (mbsf) -80 0

δD (‰, VSMOW) -70

-60

-50

depth (mbsf) -80 0

-30

-40

(c) δD (‰, VSMOW) -70

-50

-60

-30

-40

Layer 1

Layer 0

Layer 2

Layer 1

20

267

depth (mbsf) -80 0

-70

-60

δD (‰, VSMOW) -50 -40 -30 Layer 0 Layer 1

20

20 Layer 2

Layer 2

40

40

40 Layer 3

Layer 3 Layer 3

60

60

60

smectite kaolinite + illite kaolinite + smectite I/Sm

80

corrensite chlorite or Chl/Sm illite or I/Sm + chlorite or Chl/Sm

100

120 Fig. 3. Depth profiles of hydrogen isotope ratios of clay fraction samples. (a) Site C0013. (b) Site C0014 (hole B). (c) Site C0014 (hole G).

higher than those in Layer 2, except for one sample (C0014G 18T-4 2035). The value for sample C0014G 18T-4 20-35 was similar to those measured in Layer 2 (Table 1). Depth profiles of hydrogen isotope ratios of the clay fraction samples are shown in Fig. 3. Hydrogen isotope values (−40 to −55‰) did not differ significantly or systematically between the layers, except that the clay fractions containing smectite in Layer 1 at Site C0014 had relatively low values between −65 and −72‰ (Fig. 3b, c). Similarly, low hydrogen isotope ratios in clay fractions containing smectite were measured in seafloor hydrothermal fields in Kagoshima Bay (Miyoshi et al., 2013).

4.2. Chemical composition of clay fractions in Layer 3 Clay fraction samples from Layer 3 consisted of illite (or illite/smectite) and chlorite (or chlorite/smectite) (Miyoshi et al., 2015). To assign oxygen isotope values to these clay minerals, it is necessary to know their proportions. We conducted chemical analyses of the clay fraction samples in Layer 3 by the Cu-EPMA method described in Miyoshi et al. (2017) (Table 2a). We also conducted TEM-EDS analyses of individual clay mineral particles, illite (or illite/smectite) in Layer 3, and chlorite (or chlorite/ smectite) in Layers 2 and 3 (Table 2b). Results of these analyses are illustrated in Fig. 4 for the major components SiO2, Al2O3, K2O, and MgO. In

Table 2 Result of chemical analysis. (a) Chemical composition of clay fraction determined by Cu-EPMA method Sample ID

Depth (mbsf)

Minerala

N

SiO2

TiO2

Al2O3

FeO

MnO

MgO

CaO

Na2O

K2O

SO3

Total

C0013E 9X-CC 8-10 C0014B 5H-13 53-65 C0014B 5H-15 80-91 C0014G 6H-3 47-60 C0014G 9X-2 25-40 C0014G 12H-3 15-25 C0014G 14T-2 15-30 C0014G 16T-1 52-62 C0014G 18T-4 20-35 C0014G 20T-1 43-63 C0014G 23X-1 12-22 C0014G 25T-1 35-45

45.1 42.0 44.7 48.8 55.9 65.3 71.8 76.4 87.2 94.1 104.8 114.1

I, Chl I/Sm, Chl I/Sm, Chl I/Sm, Chl I/Sm, Chl I/Sm, Chl/Sm I/Sm, Chl/Sm I/Sm, Chl/Sm I, Chl I, Chl I, Chl I, Chl

13 5 5 10 10 10 10 13 20 10 20 13

45.4 50.6 53.2 51.1 49.1 48.3 48.2 51.7 39.9 50.0 53.6 50.7

0.1 0.2 0.1 0.2 0.2 0.2 0.2 0.1 0.1 b0.1 0.2 0.1

27.8 31.0 31.9 30.9 30.3 29.3 30.1 31.5 30.2 32.8 32.1 33.7

2.5 0.7 0.4 0.6 1.2 1.5 1.4 0.9 3.4 1.2 0.8 1.2

0.3 0.2 0.1 0.2 0.2 0.3 0.2 0.1 0.3 0.1 0.1 0.1

18.9 10.2 5.5 9.5 11.8 13.2 13.3 7.3 22.8 7.1 3.7 5.4

0.1 0.1 0.1 0.1 0.1 0.2 0.1 0.1 0.1 0.1 0.1 0.2

0.045 0.1 0.2 0.1 0.1 0.1 0.1 0.2 0.0 0.2 0.2 0.3

4.7 6.8 8.5 7.3 6.9 6.6 6.3 8.0 3.1 8.3 8.5 8.3

0.2 0.1 0.1 b0.1 0.1 0.3 b0.1 0.1 0.2 0.2 0.7 0.1

100 100 100 100 100 100 100 100 100 100 100 100

(b) Chemical composition of clay mineral determined by TEM-EDS Sample ID

Depth (mbsf)

C0013D 2H-2 16-18a C0013D 2H-6 14-16 C0013E 6X-CC 22-24a C0013E 7L-2 87.5-99.5 C0013E 9X-CC 8-10 C0014G 4H-5 45-62 C0014G 4H-11 80-95 C0013E 9X-CC 8-10 C0014G 6H-3 47-59 C0014G 20T-1 43-63 C0014G 25T-1 35-45

13.7 18.0 23.2 27.2 45.0 31.0 37.0 45.0 49.0 94.0 114.0

Minerala

N

SiO2

TiO2

Al2O3

Fe2O3

MgO

MnO

CaO

Na2O

K2O

SO3

Total

Chl Chl Chl Chl Chl Chl/Sm Chl/Sm I I/Sm I I

13 10 10 12 7 9 10 12 12 10 14

37.1 37.2 36.1 37.2 36.9 41.3 37.7 51.2 52.8 51.5 50.8

n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.

24.3 22.9 24.0 23.9 26.1 21.7 24.5 33.1 33.1 34.9 34.7

1.6 1.7 4.1 5.9 4.4 0.6 0.8 1.5 1.0 1.1 1.2

36.7 37.8 35.2 32.5 30.8 35.7 36.4 4.49 4.01 3.44 3.64

0.4 0.4 0.5 0.8 0.5 0.2 0.4 n.d. n.d. n.d. n.d.

n.d. n.d. 0.3 0.5 n.d. n.d. 0.2 n.d. 0.3 n.d. 0.2

n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 0.29

n.d. n.d. n.d. 0.6 1.2 0.5 0.2 9.7 8.9 9.0 9.2

n.d. n.d. 0.3 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.

100 100 101 101 100 100 100 100 100 100 100

n.d.: not detected. Abbreviations for clay minerals; I: illite, I/Sm: illite/smectite mixed-layer mineral, Chl: chlorite, Chl/Sm: chlorite/smectite mixed-layer mineral. a Miyoshi et al. (2015).

Y. Miyoshi et al. / Journal of Volcanology and Geothermal Research 384 (2019) 263–274

(1)

50

(2) 40

30

40

(1)

35

K2O (wt%)

60

Al2O3 (wt%)

SiO2 (wt%)

268

(2)

30 25

40 35 30 25

20

20

15

15

10

10

5

5

20

(1)

10

0

0

0 0

10

20

30

40

50

60

(2)

0

5

10

15

20

MgO (wt%)

25

30

35

40

0

5

10

15

20

25

30

35

40

MgO (wt%)

MgO (wt%)

Clay fraction sample in Layer 3 analyzed by Cu-EPMA method (Miyoshi et al., 2017) Illite or illite/smectite mixed-layer mineral in Layer 3 analyzed by TEM-EDS Chlorite or chlorite/smectite mixed-layer mineral in Layers 2 and 3 analyzed by TEM-EDS (1) C0013E 9X-CC 8-10, (2) C0014G 18T-4 20-35 Fig. 4. SiO2, Al2O3, K2O, and MgO contents in clay fraction samples in layer 3.

each diagram, plots of the clay fraction samples are aligned along a mixing line between those of illite and chlorite minerals. Two samples (C0013E 9X-CC 8-10 and C0014G 18T-4 20-35) exhibited higher proportions of chlorite, while other samples had closer resemblance to illite. We estimated the proportion of illite (and illite/smectite) in the clay fraction samples in Layer 3 based on their K2O content (Table 3). Because K is contained in illite and illite/smectite as an interlayer cation, but rarely contained in chlorite and chlorite/smectite, K2O contents in the clay fraction samples should be correlated with the proportion of illite. We assumed that the K2O content of illite (or illite/smectite) in the clay fraction samples in Layer 3 was equal to that of illite in the sample collected from the maximum depth at each site to simplify the calculation (Table 3). In fact, the K2O contents of illite (and illite/smectite) determined by TEM-EDS showed good agreement for three samples collected from hole C0014G (Table 2b). The proportion of illite was calculated as 69–93% for the clay fraction samples in Layer 3 at site C0014, except for sample C0014G 18T-4 20-35 in which the proportion of illite was estimated to be 34% (Table 3). The proportion of illite was calculated as 52% for sample C0013E 9X-CC 8-10, which indicates that the proportion of chlorite was comparable to that of illite. The observed variation in illite proportions in Layer 3 is in agreement with the variations in mineralogical composition of sediment cores from the same site, previously determined by XRD analysis (Shao et al., 2017).

4.3. Isotope equilibria between the fluid and clay minerals We calculated formation temperatures of the clay minerals based on the temperature dependence of oxygen and hydrogen isotope fractionation between the fluid and clay minerals. Isotopic equilibrium between clay minerals and fluid is assumed during clay mineral formation because the previous studies revealed that occurrence of the clay minerals suggests the formation by intense and uniform hydrothermal alteration (Miyoshi et al., 2015; Shao et al., 2017). Since the study site is an active hydrothermal field, it is reasonable that the oxygen and hydrogen isotope ratios of the fluid can be deduced from those of pore fluids collected from the sediment cores. Chemical analyses of these pore fluids showed that they were composed of a hydrothermal and a seawater component (Takai et al., 2011). Discharge of the hydrothermal fluid from some holes drilled at sites C0013, C0014, and C0016 was documented in previous studies (Takai et al., 2011; Kawagucci et al., 2013). Furthermore, Kawagucci et al. (2013) reported oxygen and hydrogen isotope ratios of the hydrothermal fluids collected from these holes. The oxygen isotope values were +1.2‰ at site C0016, +1.1‰ at site C0013, and +1.3‰ at site C0014 and for hydrogen isotopes they were −1.0‰ at site C0016, −1.9‰ at site C0013, and −0.8‰ at site C0014 (Kawagucci et al., 2013). We calculated the average value of the oxygen and hydrogen isotope ratios reported in

Table 3 Illite proportion of the clay fraction samples in Layer 3. Sample ID

C0013E 9X-CC 8-10 C0014B 5H-13 53-65 C0014B 5H-15 80-91 C0014G 6H-3 47-59 C0014G 9X-2 25-40 C0014G 12H-3 15-25 C0014G 14T-2 15-30 C0014G 16T-1 52-62 C0014G 18T-4 20-35 C0014G 20T-1 43-63 C0014G 23X-1 12-22 C0014G 25T-1 35-45

Depth (mbsf)

45.1 42.0 44.7 48.8 55.9 67.3 71.8 81.3 87.2 94.1 104.8 114.1

Dominant clay minerala

I, Chl I/Sm, Chl I/Sm, Chl I/Sm, Chl I/Sm, Chl I/Sm, Chl/Sm I/Sm, Chl/Sm I/Sm, Chl/Sm I, Chl I, Chl I, Chl I, Chl

K2O content (wt%)

Illite proportion (%)

Clay fraction

Illite

4.7 6.8 8.5 7.3 6.9 6.6 6.3 8.0 3.1 8.3 8.5 8.3

9.71 9.22 9.22 9.22 9.22 9.22 9.22 9.22 9.22 9.22 9.22 9.22

Abbreviations for clay minerals; I: illite, I/Sm: illite/smectite mixed-layer mineral, Chl: chlorite, Chl/Sm: chlorite/smectite mixed-layer mineral. a Miyoshi et al. (2015). 1 K2O content of illite in C0013E 9X-CC 9-10. 2 K2O content of illite in C0014G 25T-1 35-45.

52 74 93 80 75 72 69 87 34 90 93 90

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Kawagucci et al. (2013) and considered it to be that of the hydrothermal component (+1.2 ± 0.1‰ and −1.2 ± 0.5‰, respectively). Considering mixing between the hydrothermal and seawater components, a range between ~0 and +1.2‰ and between ~0 and −1.2‰ is assumed for the oxygen and hydrogen isotope ratios of the fluid. 4.4. Calculation of formation temperatures of the clay minerals 4.4.1. Based on oxygen isotope equilibrium Smectite, kaolinite, and illite/smectite were determined to be the dominant clay minerals in Layer 1. Using data from the clay fraction samples that consisted exclusively of smectite, oxygen isotope equilibrium temperatures were calculated using the fractionation equation from Sheppard and Gilg (1996). We did not consider effects of the chemical composition of smectite on the isotope fractionation factor, in spite of the discussion surrounding this in a previous study (Savin and Lee, 1988). Sheppard and Gilg (1996) mentioned that it is difficult to determine the effect on oxygen isotope fractionation precisely, because of the paucity of data available. Results of this calculation are listed as estimated formation temperatures in Table 4a. The estimated temperature range for the formation of smectite was 111–175 °C. The dominant clay minerals in Layer 2 were chlorite, chlorite/smectite, and corrensite. The chlorite/smectite in Layer 2 consisted almost

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entirely of chlorite based on XRD analysis (Miyoshi et al., 2015), so we apply the oxygen isotope fractionation equation between chlorite and fluid to these samples. Another factor that must be considered is the effect of the chemical composition of chlorite on oxygen isotope fractionation, which has been discussed in previous studies (e.g., Savin and Lee, 1988; Cole and Ripley, 1999). Lacroix and Vennemann (2015) suggested that compositional effects for oxygen isotope fractionation are not important for chlorites that have Fe / (Fe + Mg) between 0.35 and 0.7, but are significant for Mg-rich chlorites. The chlorite and chlorite/ smectite in our study site are Mg-rich (Fe / (Fe + Mg) b 0.1) (Shao et al., 2015; Miyoshi et al., 2015; Yeats et al., 2017). Following the discussion of Lacroix and Vennemann (2015), we calculated oxygen isotope equilibrium temperatures of chlorite and chlorite/smectite by applying the formula proposed by Wenner and Taylor (1971), which discussed highly Mg-rich chlorites associated with serpentine in hydrothermally altered rocks. The calculated results are listed in Table 4b, with an estimated temperature range for the formation of chlorite or chlorite/smectite from 177 to 297 °C. Calculation of the oxygen isotope equilibrium temperatures for clay minerals in Layer 3 requires another step, because the samples in Layer 3 consisted of two clay minerals: chlorite (or chlorite/smectite) and illite (or illite/smectite). As discussed previously in Section 4.2, chemical analyses revealed that the two clay minerals were dominated by illite,

Table 4 Formation temperature of clay minerals estimated based on their oxygen isotope ratios. (a) Smectite in Layer 1 Sample ID

C0013C 1H-3 22-24 C0014B 3H-2 52-62 C0014B 3H-5 20-30 C0014G 3H-2 50-65 C0014G 3H-5 32-47

Depth (mbsf)

4.7 17.4 20.1 19.8 22.9

Clay mineral for estimation

Smectite Smectite Smectite Smectite Smectite

Formation temperature (°C) T (s)

T (h)

133 129 155 111 130

150 146 175 125 147

(b) Chlorite or chlorite/smectite mixed-layer mineral in Layer 2 Sample ID

C0013D 2H-2 16-18 C0013D 2H-6 14-16 C0013E 6X-CC 22-24 C0013E 7L-2 87.5-99.5 C0014B 4H-2 35-45 C0014B 4H-4 35-45 C0014B 4H-7 111-121 C0014G 4H-5 45-60 C0014G 4H-7 0-15 C0014G 4H-10 0-15 C0014G 4H-11 80-95

Depth (mbsf)

13.7 18.0 23.2 27.2 26.4 29.2 33.0 30.9 32.4 35.5 37.3

Clay mineral for estimation

Chlorite Chlorite Chlorite Chlorite Chl/Sm Chl/Sm Chl/Sm Chl/Sm Chl/Sm Chl/Sm Chl/Sm

Formation temperature (°C) T (s)

T (h)

225 196 196 177 181 209 207 180 237 190 227

281 242 242 217 222 260 256 221 297 233 284

(c) Illite or illite/smectite mixed-layer mineral in Layer 3 Sample ID

C0013E 9X-CC 8-10 C0014B 5H-13 53-65 C0014B 5H-15 80-91 C0014G 6H-3 47-59 C0014G 9X-2 25-40 C0014G 12H-3 15-25 C0014G 14T-2 15-30 C0014G 16T-1 52-62 C0014G 18T-4 20-35 C0014G 20T-1 43-63 C0014G 23X-1 12-24 C0014G 25T-1 35-45

Depth (mbsf)

45.1 42.0 44.7 48.8 55.9 65.3 71.8 76.4 87.2 94.1 104.8 114.1

Clay mineral for estimation

Illite I/Sm I/Sm I/Sm I/Sm I/Sm I/Sm I/Sm Illite Illite Illite Illite

Formation temperature (°C) T (s)

T (h)

(309) 220 238 241 241 238 256 268 (293) 262 265 262

(367) 253 276 279 279 279 298 314 (346) 306 310 306

T(s): Formation temperature based on oxygen isotope equilibrium between clay mineral and seawater (±0 ‰). T(h): Formation temperature based on oxygen isotope equilibrium between clay mineral and hydrothermal fluid (+1.2‰) (See text for details). The temperatures in parentheses were excluded in discussion because the estimation for these temperatures included too much uncertainties (see text for details). Abbreviations for clay minerals; Chl/Sm: chlorite/smectite mixed-layer mineral, I/Sm: illite/smectite mixed-layer mineral.

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with the exception of two samples, namely C0013E 9X-CC 8-10 and C0014G 18T-4 20-35 (Fig. 4, Table 3). Using the proportion of the two clay minerals, based on chemical analyses, we estimate formation temperature of the dominant illite (or illite/smectite). Considering the narrow range of the oxygen isotope values of chlorite (and chlorite/ smectite) in Layer 2 (from +1.3 to +3.0‰; Table 1), it would be reasonable to assume that the value of the minor chlorite (and chlorite/smectite) in Layer 3 is not too different from that of the deepest sample in Layer 2. The assigned oxygen isotope values for illite are listed in Table 5, together with the assumed oxygen isotope values for chlorite. The oxygen isotope values of illite shift by +0.3–1.4‰ from the measured oxygen isotope values of the clay fraction samples. Oxygen isotope equilibration temperatures between illite and water were calculated according to the temperature dependence proposed by Sheppard and Gilg (1996). Although the temperature dependence of the oxygen isotope equilibrium between illite/smectite and water was proposed by Savin and Lee (1988), we did not consider illite/smectite, because XRD analysis indicated that the illite/smectite in Layer 3 consisted mostly of illite layers (Miyoshi et al., 2015). Results of the calculation are listed in Table 4c. Oxygen isotope equilibrium temperatures are 20–40 °C higher, if we neglect the chlorite oxygen contribution, which is unreasonable considering the thermal stability of clay minerals (e.g., Inoue, 1995). The calculated formation temperatures for the samples C0013E 9X-CC 8-10 and C0014G 18T-4 20-35 are also higher than would be expected for a reasonable range. These estimations have high uncertainties, because these samples consisted of high proportions of chlorite, for which the oxygen isotope values were assumed. We thus exclude the estimated formation temperatures for these two samples in further discussions. 4.4.2. Based on hydrogen isotope equilibrium In a reinterpretation of published data, Gilg and Sheppard (1996) suggested that the hydrogen isotope fractionation factor between kaolinite and water varied monotonically, however, the variation is relatively small (~5‰ over the temperature of 300 to 100 °C). This was in contrast to an earlier study by Lambert and Epstein (1980) who indicated a much more pronounced and non-linear fractionation between kaolinite and water. In the case of smectite, the hydrogen-water isotope fractionation relative to temperature (between 300 and 100 °C) shows no consistent variation, there is no data for temperatures between 100 and 25 °C and a variation of almost 60‰ for temperatures between 25 and 10 °C (Fig. 4 in Gilg and Sheppard, 1996). Capuano (1992) and Yeh (1980) reported a temperature dependence of hydrogen isotope equilibrium for illite-water and smectite-water, however, this was

over a low temperature range (0–150 °C and 0–120 °C, respectively). Graham et al. (1987) reported hydrogen isotope fractionation factors for chlorite and water, but over a limited range and at relatively high temperatures (400–500 °C). So, any calculated equilibrium isotope temperatures for chlorite presume that the curve is linear and can be extended to lower temperatures. In contradiction to all these studies, Marumo et al. (1980) concluded that in the temperature range from 100 to 250 °C, hydrogen isotopic fractionation factors between clay minerals and water are not sensitive to the temperature. Assuming the Gilg and Sheppard (1996) hydrogen isotope fraction factors to be correct, the calculated isotope equilibrium temperature from kaolinite (about −40‰) and water (−1.2‰, Kawagucci et al., 2013) is −9 °C. If the Lambert and Epstein (1980) hydrogen isotope fraction factors were assumed to be correct, the calculated isotope equilibrium temperature for kaolinite and water is +5 °C. Calculated equilibrium temperatures for smectite ranged from +35 to −74 °C and for illite from +48 to −52 °C (Capuano, 1992 and Yeh, 1980), with the majority of the calculated temperatures being negative. The equilibrium isotope temperatures calculated using the chlorite-water fraction curve from Graham et al. (1987) average +145 °C (range 112–242 °C), which is higher than temperatures calculated from other clay minerals for hydrogen isotope fractionation, but still lower than temperatures calculated from the oxygen isotope results (177–297 °C). The calculated hydrogen isotope equilibrium temperatures using published fractionation factors of clay minerals and water provide unrealistic and inconsistent data for the various clay minerals. So, in this study we have not considered equilibrium temperature calculations based on hydrogen isotope data from clay minerals and agree with Savin and Hsieh (1998) who said that the equilibrium hydrogen isotope fractionation between clay minerals is complex and not well established and that additional work is required. 4.5. Evaluation of the calculated formation temperatures from oxygen isotope data The estimated formation temperature ranges for each clay mineral in this study are summarized in Table 6, which gives a comparison to those of other seafloor hydrothermal fields reported in previous studies. The estimated temperature range for the formation of smectite was 111–175 °C. This range is comparable to the temperature range reported in previous studies, with the exception of two studies that proposed higher temperatures of up to 240–250 °C (Table 6). For on-land geothermal fields, the formation temperature of smectite is generally considered to be b100 °C (e.g., Inoue, 1995), however, smectite in

Table 5 Estimated oxygen isotope ratios of illite or illite/smectite mixed-layer mineral in Layer 3. Sample ID C0013E 9X-CC 8-10 C0014B 5H-13 53-65 C0014B 5H-15 80-91 C0014G 6H-3 47-59 C0014G 9X-2 25-40 C0014G 12H-3 15-25 C0014G 14T-2 15-30 C0014G 16T-1 52-62 C0014G 18T-4 20-35 C0014G 20T-1 43-63 C0014G 23X-1 12-24 C0014G 25T-1 35-45

Depth (mbsf) 45.1 42.0 44.7 48.8 55.9 65.3 71.8 76.4 87.2 94.1 104.8 114.1

Dominant clay minerala Illite, chlorite I/Sm, chlorite I/Sm, chlorite I/Sm, chlorite I/Sm, chlorite I/Sm, Chl/Sm I/Sm, Chl/Sm I/Sm, Chl/Sm Illite, chlorite Illite, chlorite Illite, chlorite Illite, chlorite

δ18O (clay fraction) (‰) 3.1 4.7 5.1 4.3 4.0 4.0 3.5 3.8 2.1 4.2 4.2 4.2

δ18O (chlorite) (‰) 1

3.0 2.12 2.12 1.53 1.53 1.53 1.53 1.53 1.53 1.53 1.53 1.53

δ18O (illite) (‰) (3.3) 6.1 5.4 5.3 5.3 5.4 4.8 4.4 (3.7) 4.6 4.5 4.6

δ18O (clay fraction): Oxygen isotope ratio for the clay fraction sample, δ18O (chlorite): Oxygen isotope ratio for chlorite or chlorite/smectite mixed-layer mineral, δ18O (illite): Oxygen isotope ratio for illite or illite/smectite mixed-layer mineral. Abbreviations for clay minerals; I/Sm: illite/smectite mixed-layer mineral, Chl/Sm: chlorite/smectite mixed-layer mineral. The δ18O (illite) in parentheses were excluded in discussion because the estimation for these values included too much uncertainties (see text for details). a Miyoshi et al. (2015). 1 Oxygen isotope ratio of chlorite in C0013E 7L-2 87.5-99.5. 2 Oxygen isotope ratio of chlorite/smectite mixed-layer mineral in C0014B 4H-7 111-121. 3 Oxygen isotope ratio of chlorite/smectite mixed-layer mineral in C0014G 4H-11 80-95.

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Table 6 Comparison of estimated formation temperatures of clay minerals in seafloor hydrothermal fields. Clay mineral Smectite (montmorillonite) Mg-Fe-smectite Montmorillonite Montmorillonite Mg-smectite (saponite) Montmorillonite Chlorite or chlorite/smectite Chlorite Chlorite Chlorite Chlorite Illite or illite/smectite Illite (mica) Illite

Estimated formation temperature [°C] Formulaa

Depth [mbsf] Location ~25 ~1 ~3 ~4 ~4 ~10 7–114 ~2 16–266 144–239 116–197 28–114 Surface 40–355

Iheya-North (1) Escanaba (2) Grimsey (3) Wakamiko (4) Wakamiko (4) Suiyo (5) Iheya-North (1) Escanaba (2) Middle Valley (AAV) (6) Middle Valley (BH) (6) PACMANUS (7) Iheya-North (1) Jade (8) PACMANUS (7)

111–175 181–240 95–116 118–163 164 166–251 177–297 199–221 125–280 180–320 220–265 220–314 238, 145 250–310

Reference

Sheppard and Gilg (1996) Savin and Lee (1988) Savin and Lee (1988) Sheppard and Gilg (1996) Sheppard and Gilg (1996) Yeh and Savin (1977) Wenner and Taylor (1971) Wenner and Taylor (1971) Savin and Lee (1988) Wenner and Taylor (1971) Wenner and Taylor (1971) Sheppard and Gilg (1996) Sheppard and Gilg (1996) Sheppard and Gilg (1996)

This study Zierenberg and Shanks III (1994) Dekov et al. (2008) Miyoshi et al. (2013) Miyoshi et al. (2013) Marumo et al. (2008) This study Zierenberg and Shanks III (1994) Buatier et al. (1995) Lackschewitz et al. (2000) Lackschewitz et al. (2004) This study Marumo and Hattori (1999) Lackschewitz et al. (2004)

Abbreviations for clay minerals; illite/smectite: illite/smectite mixed-layer mineral, chlorite/smectite: chlorite/smectite mixed-layer mineral. (1) Iheya-North-Knoll, Okinawa Trough, (2) Escanaba Trough, Gorda Ridge, (3) Grimsey Graben, (4) Wakamiko crater, Kagoshima Bay, (5) Suiyo seamount, Izu Bonin-Arc, (6) Middle Valley, Juan de Fuca Ridge, (AAV) the Area of Active Venting, (BH) the Bent Hill sediment mound, (7) PACMANUS, Manus back-arc basin, (8) Jade, Izena Hole, Okinawa Trough. a Formula of temperature dependence of oxygen isotope equilibrium between clay mineral and water.

seafloor hydrothermal systems appears to be more stable at higher temperatures. The estimated temperature range for chlorite and chlorite/smectite formation was 177–297 °C, which is in general agreement with those reported for other seafloor hydrothermal systems (Table 6). The estimated temperature range for illite and illite/smectite formation was 220–314 °C, which is also in general agreement with those reported in the previous studies (Table 6). The estimated temperatures for chlorite and illite are also comparable to approximate temperature ranges for

(a) depth (mbsf) 0

chlorite and illite formation (200–300 °C) in on-land geothermal fields (e.g., Inoue, 1995). The temperature of hydrothermal fluids discharging from natural vents at the NBC mound and drill holes was reported to range from 240 °C to 311 °C (Nakagawa et al., 2005; Takai et al., 2011; Kawagucci et al., 2013), which is close to the boiling point of water at the seafloor depth (~1000 m). The estimated formation temperature range for chlorite (or chlorite/smectite) and illite (or illite/smectite) is also comparable to the hydrothermal fluid temperature.

(b) Temperature ( oC) 0 Layer 1

100

200

Temperature ( oC)

300

0

4 ~ 6 oC/m 20

12 mbsf

Layer 1

Layer 2 28 mbsf

300

6 ~ 9 C/m 16 mbsf

25 mbsf

Layer 2 40 mbsf

40 Layer 3

200 o

Layer 0

6 mbsf

100

Layer 3

37.6 mbsf 47 mbsf 50 mbsf

60

80

100

120 Formation temperature of clay mineral estimated by our study Formation temperature of smectite Formation temperature of chlorite or chlorite/smectite mixed-layer mineral Formation temperature of illite or illite/smectite mixed-layer mineral

Downhole temperature measured during the expedition (Takai et al., 2011) Average bottom water temperature Measured using advanced piston corer temperature tool (APCT3) Measured using thermoseal temperature-sensitive strips Estimated temperature range

Fig. 5. Depth profiles of the estimated formation temperatures of clay minerals. (a) Site C0013; (b) Site C0014. The left end of the temperature bars indicate the formation temperature based on oxygen isotope equilibrium between the clay mineral and seawater assuming a δ18O value of ~0‰ and the right ends indicate those based on oxygen isotope equilibrium between the clay mineral and hydrothermal fluid assuming a δ18O value of +1.2‰ (see text for details). Oblique thick gray lines indicate thermal gradient at each Site.

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4.6. Subseafloor thermal gradient below a hydrothermal field at the Iheya North Knoll Estimated formation temperatures of the clay minerals are plotted against depth in Fig. 5. At site C0013, located about 100 m east of the NBC mound, a temperature of N100 °C was determined for smectite at a shallow depth of ~6 mbsf (Layer 1), and ~200 °C was determined for chlorite at depths from ~6 to ~28 mbsf (Layer 2) (Fig. 5a). At site C0014, located about 450 m east of the NBC mound, a temperature of N100 °C was determined for smectite at shallow depths from ~12 mbsf to ~25 mbsf (Layer 1), ~200 °C was determined for chlorite at depths from ~25 to ~40 mbsf (Layer 2), and N220 °C was determined for illite below 40 mbsf (Layer 3) (Fig. 5b). During IODP Expedition 331, borehole temperature information was obtained using two devices at site C0014 (Takai et al., 2011). The results of these measurements are plotted against depth in Fig. 5b for comparison to the oxygen isotope equilibrium calculated formation temperatures for clay minerals. Temperatures below 55 °C were obtained by in situ measurement using a temperature probe (APCT3), which was kept in the sediment for N5 min to allow the measured temperature to stabilize. These results are plotted as circles in Fig. 5b; three measurements in Layer 0 showed temperatures of 16–23 °C, and the temperature at 16 mbsf exceeded 55 °C, which was the maximum calibrated temperature of the probe. The range of measured temperature in Layer 0 is in accordance with the finding that altered clay minerals were not observed in this layer. The fact that in situ measurements failed at 16 mbsf in Layer 1 due to the high temperature is also in accordance with the clay mineral assemblage and estimated formation temperatures for Layer 1, even though no exact temperature information was available. Borehole temperature information at deeper depths was obtained using thermoseal temperature-sensitive strips taped to the outer surface of the core liner (Takai et al., 2011, 2012). These strips have embedded, chemically impregnated wafers that turn black when exposed to a designated temperature (crosses in Fig. 5b). Two of the

three temperature measurements obtained using thermoseal strips exceeded the range of all the beads on the strip used and are thus estimated lower limits of temperature. The borehole temperature was estimated as N120 °C at 37.6 mbsf and N210 °C at 50.2 mbsf. The borehole temperature was reported to be 145 ± 5 °C at 47 mbsf, because the beads in the thermoseal strips indicated that the temperature reached or exceeded 140 °C, but did not reach 150 °C. These borehole temperatures are lower than the range of estimated formation temperatures for clay minerals in the corresponding depths, as illustrated in Fig. 5b. This discrepancy might indicate that the present temperature gradient was lower than in the period when the clay minerals were formed by hydrothermal activity. However, it is difficult to estimate the true discrepancy because of ambiguity in the temperature information obtained using thermoseal strips. It is notable that the thermal gradient below the seafloor was significant, since temperatures of N200 °C were determined at 50 mbsf in both estimations. The thermal gradient is calculated to be 6–9 °C/m for Site C0014 between the seafloor and Layer 2 (Fig. 5b). For Site C0013, the thermal gradient between Layers 1 and 2 is 4–6 °C/m, although extrapolation of the gradient leads to an unreasonable high temperature at the seafloor (Fig. 5a). The thermal gradient appears to be b1 °C/m within Layer 3 at Site C0014. This highly variable thermal gradient could be explained by the existence of a heat source at shallow depth. Takai et al. (2011) noted that after drilling, fluid discharged that had a temperature higher than 240 °C at Site C0013 and C0014. Within the seafloor hydrothermal system, a high temperature fluid could be stable just below the seafloor, because the pressure prevents boiling. Considering that the seafloor is always cooled by cold (1–5 °C) seawater, it is not surprising that a steep thermal gradient is observed in the subseafloor region. The estimated formation temperature ranges obtained in this study are in good agreement with the layered occurrence of distinctive assemblages of clay minerals in the hydrothermal field at Iheya North Knoll revealed by a previous study (Miyoshi et al., 2015), as illustrated in Fig. 6. This agreement suggests that the occurrence of clay minerals generally

Fig. 6. Subseafloor temperature gradient in the Iheya North Knoll based on the oxygen isotope equilibrium temperatures of clay minerals. The layered occurrence of clay minerals was revealed in a previous study (Fig. 10 in Miyoshi et al., 2015). The temperature in parentheses indicates the borehole temperature measured during the expedition (Takai et al., 2011).

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reflects the subseafloor thermal gradient below the hydrothermal field. Furthermore, the profile of the estimated formation temperatures does not show a uniform increase with depth, but rather a stepwise increase. In Layer 0 of the unaltered layer, the in situ measured temperatures at site C0014 revealed constant temperatures of 16–23 °C, while the estimated temperature in underlying Layer 1 exceeded 111 °C. Another temperature gap was found between Layer 1 and Layers 2 and 3: 111–175 °C in Layer 1, and 177–314 °C in Layers 2 and 3. As already discussed in Miyoshi et al. (2015), impermeable hard sediment layers are taken to be responsible for the layers of distinctive assemblages of clay minerals, by bounding the hydrological structure beneath the hydrothermal field. The temperature profiles obtained in this study likely reflect the hydrological structure bounded by the impermeable hard sediment. 5. Summary Occurrence of diverse clay minerals in the sediment beneath the active hydrothermal field in the Iheya North Knoll, Okinawa Trough had been revealed by our previous study of deep sediments collected during IODP Expedition 331 (Miyoshi et al., 2015). The sediments were classified into four layers, characterized by distinctive clay mineral assemblages. Some of the clay fractions of the sediments were revealed to be composed exclusively of one or two clay minerals. We attempted to estimate the formation temperatures of the clay minerals from their oxygen and hydrogen isotope ratios, assuming that they were isotopically equilibrated with surrounding hydrothermal fluids. The oxygen and hydrogen isotope values of the fluids were estimated to range from ~0 to +1.2‰ and ~0 to −1.2‰, respectively, based on measured isotope values of the seawater and vent fluid. The calculated oxygen isotope temperatures ranged from 111 to 175 °C for Layer 1, which contained smectite and a minor kaolinite component, 177 to 297 °C for Layer 2, which was dominated by chlorite or chlorite/smectite, and 220 to 314 °C for Layer 3, which contained illite or illite/smectite (with less chlorite). The temperature ranges obtained are in general agreement with the formation temperatures of each clay mineral in other seafloor hydrothermal fields discussed in previous studies. The depth profile of the formation temperatures was not completely confirmed by the results of downhole temperature measurements during the same expedition. However, available measurements were consistent in that they suggested a significantly steep, step-wise, thermal gradient below the active hydrothermal field, where temperatures of N200 °C were determined at depths as low as 50 mbsf. Stable isotope studies of the clay minerals provided quantitative information that enabled us to draw a profile of the subseafloor thermal gradient of the hydrothermal system at Iheya North Knoll. Temperatures calculated from hydrogen isotope data, based on current published hydrogen isotope fractionation data between clay minerals and water, gave unrealistic (negative) and inconsistent temperatures. Additional work on hydrogen isotope fractionation between clay minerals and water is still required. Acknowledgments We are grateful to the IODP, captains, scientific party, expedition staff, onboard technicians, and crew who assisted with drilling, sampling, and measurements during IODP Expedition 331. We are grateful to Mr. Andy Phillips and Ms. Jannie Cooper from GNS Science for their assistance with stable isotope analyses. We thank the staff of the Research Laboratory for HVEM at Kyushu University for their support with TEM analyses. We appreciate an anonymous reviewer for his/her instructive comments that helped to improve significantly earlier version of the manuscript, and editor Prof. Alessandro Aiuppa for handling this paper. This manuscript forms part of the doctoral thesis of the first author (Y. M.). This study was partly supported by the “TAIGA project”, which was funded by a Grant-in-Aid for Scientific Research on

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