Mg-rich clay mineral formation associated with marine shallow-water hydrothermal activity in an arc volcanic caldera setting

Chemical Geology 355 (2013) 28–44

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Mg-rich clay mineral formation associated with marine shallow-water hydrothermal activity in an arc volcanic caldera setting Youko Miyoshi a,⁎, Jun-ichiro Ishibashi a, Kevin Faure b, Kotaro Maeto c, Seiya Matsukura a, Akiko Omura d, Kazuhiko Shimada a, Hiroshi Sato e, Takeaki Sakamoto a, Seiichiro Uehara a, Hitoshi Chiba c, Toshiro Yamanaka c a

Department of Earth and Planetary Sciences, Graduate School of Sciences, Kyushu University, Japan Division of Geological Resources, GNS Science, Lower Hutt, New Zealand c Department of Earth Sciences, Graduate School of Natural Science and Technology, Okayama University, Japan d Graduate School of Frontier Science, The University of Tokyo, Japan e School of Business Administration, Senshu University, Japan b

a r t i c l e

i n f o

Article history: Received 18 April 2013 Received in revised form 28 May 2013 Accepted 28 May 2013 Available online 6 June 2013 Editor: J. Fein Keywords: Hydrothermal alteration Montmorillonite Saponite Kerolite Pore fluid chemistry Isotope

a b s t r a c t Shallow-water hydrothermal activity, represented by venting of hydrothermal fluid around 200 °C, occurs in the Wakamiko submarine crater at 200 m water depth in Kagoshima Bay, southwest Japan. The crater is the center of large eruptions that formed a volcanic caldera, which is semi-submerged at present. The crater is covered with thick volcanic sediments of felsic composition. We studied the distribution and chemical composition of hydrothermal clay minerals that are abundant in the sediment collected by piston coring. We also conducted chemical analysis of pore fluids squeezed from the sediment to understand hydrothermal interactions that resulted in formation of these clay minerals. The PC-2 core (340 cm in length) collected in the vicinity of a high-temperature fluid venting site was characterized by abundant Mg-saponite that is limited to a layer between 270 and 300 centimeters below the seafloor (cmbsf) and montmorillonite throughout the core below 55 cmbsf. Vertical profiles of pore fluid chemistry suggest that saponite formation is related to the interface between the seawater and the hydrothermal component in the sediment layer. Formation temperatures of the montmorillonite were estimated to be 118–163 °C, based on oxygen isotope thermometry. Formation of the montmorillonite is attributed to hydrothermal interaction between seawater-dominant pore fluid and volcanic glass. The formation temperature of the saponite was estimated to be ~ 164 °C, based on oxygen isotope thermometry. Formation of the saponite is attributed to hydrothermal interaction between seawater-dominant pore fluid and the montmorillonite, which had been formed at a prior stage. The PC-1 core (240 cm in length) collected from a relatively low-temperature fluid shimmering site was characterized by the occurrence of kerolite in the lower section (210–240 cmbsf). Vertical profiles of pore fluid chemistry suggest that the kerolite formation occurred at the interface between seawater and the hydrothermal component of the sediment layer. Formation temperature of the kerolite was estimated to be about ~211 °C, based on oxygen isotope thermometry. Formation of the kerolite is attributed to precipitation from a fluid that was a mixture of a hydrothermal component and seawater. This study revealed the occurrence of Mg-rich clay minerals, saponite and kerolite, beneath a submarine hydrothermal field that developed within sediment of felsic composition. During hydrothermal interactions that formed these clay minerals, seawater penetrated into the sediment and was an important Mg source. Formation of Mg-rich clay minerals, saponite and kerolite, are controlled by pore fluid chemistry, which varies from a seawaterdominant to hydrothermal-dominant component. Exclusive formation of Mg-rich clay minerals at different sites could be explained by different water–rock ratios of the hydrothermal interaction — saponite formation at low water-rock ratio and kerolite precipitation at high water–rock ratio. Occurrence of Mg-rich clay minerals provides clues to the hydrological structure in sediment-covered hydrothermal systems in an arc volcanic caldera setting. © 2013 Elsevier B.V. All rights reserved.

1. Introduction ⁎ Corresponding author at: Institute for Geo-Resources and Environment, National institute of Advanced Industrial Science and Technology (AIST), Japan. E-mail address: [email protected] (Y. Miyoshi). 0009-2541/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.chemgeo.2013.05.033

Detailed investigation of clay minerals in hydrothermal areas has been considered to provide fundamental information for understanding the physicochemical conditions within the hydrothermal system. The

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

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formation of hydrothermal alteration minerals reflects several factors of fluid–mineral interactions, such as temperature, chemistry of the fluid and host-rock, and water–rock ratio. For seafloor hydrothermal systems, in general, studies of hydrothermal clays have been focused on basalt-hosted fields located at midocean ridge settings. In such systems, Mg-rich hydrothermal clays are often observed as dominant alteration minerals. At Middle Valley in the Juan de Fuca Ridge where high-temperature fluid (T = 257–276 °C) venting occurs, saponite, chlorite–smectite mixed-layer, corrensite, and chlorite were observed (Goodfellow et al., 1993; Buatier et al., 1995; Lackschewitz, 2000). Dekov et al. (2008a) reported Mg-rich smectite, chlorite, and kerolite in a shallow-water hydrothermal field at Grimsey Island north of Iceland where high-temperature fluid (T = 250 °C) venting occurs. The occurrence of Mg-rich clay minerals was also reported for a hydrothermal system in a volcanic rift setting. In the JADE hydrothermal field in the Okinawa Trough, where hydrothermal activity is represented by venting of a high-temperature fluid (T = 310 °C), Mg-rich chlorite, chlorite–smectite mixed-layer, and talc were observed, as well as Al-rich hydrothermal minerals, such as mica and kaolinite (Marumo and Hattori, 1999). Giorgetti et al. (2006) and Monecke et al. (2007) observed formation of saponite as well as montmorillonite in cracks and vesicles of hydrothermally altered dacitic lava collected from PACMANUS hydrothermal field in the Manus Basin. All of these studies considered the formation of Mg-clay minerals in seafloor hydrothermal systems to be related to mixing between the seawater and hydrothermal fluid, because seawater is the most plausible Mg source. Moreover, previous studies have demonstrated extensive fluid– sediment interaction within submarine sediment-rich hydrothermal systems (Goodfellow et al., 1993; Buatier et al., 1995; Marumo and Hattori, 1999; Lackschewitz, 2000; Dekov et al., 2008a). Compared to sediment-starved hydrothermal systems, the pore spaces of unconsolidated sediment may provide a large surface area for hydrothermal interaction. A submarine caldera of an island arc volcano would be another important geological setting for a sediment-rich hydrothermal system, since the caldera floor is often covered with a sequence of volcaniclasitc sediments produced by felsic magma activity. In this paper, we report hydrothermal clay minerals identified in unconsolidated volcanic sediment at the Wakamiko submarine crater at 200 m water depth in Kagoshima Bay, Japan. The Wakamiko crater was the center of very large eruptions related to felsic magma activity during the late Pleistocene. The eruptions formed a volcanic caldera that is semi-submerged at present. Hydrothermal activity, represented by venting of hydrothermal fluid around 200 °C, occurs on the seafloor which is thickly covered with volcanic sediment of felsic composition (Yamanaka et al., 2013). The hydrothermal system in the Wakamiko crater provides an excellent opportunity to investigate hydrothermal interaction with volcanic sediment of felsic composition. Previous studies have reported the occurrence of montmorillonite and illite–smectite mixed-layer in surface sediment (Nakaseama et al., 2008). In this study, we collected sediment from a few meters below the seafloor using a piston corer. Our goal was to characterize hydrothermal clay minerals by mineralogical, geochemical and isotopic analyses. Furthermore, we conducted geochemical analysis of the pore fluid obtained from the same core sample. Integration of results of the solid and fluid samples is used to provide constraints on the formation process of hydrothermal alteration minerals.

major volcanic activity that formed the Aira caldera during the late Pleistocene (Aramaki, 1984). Subaerially around the Kagoshima Bay and south Kyushu, a significant amount of the Ito pyroclastic flow depositions (called “shirasu” in Japanese) that erupted during the caldera-formed eruption are recognized in several places (Oba et al., 1967). Tephra from the same eruption is widely spread across Japan and recognized as “AT (Aira-Tanzawa) tephra”, which is used as a marker horizon. The main part of the Aira caldera was occupied by seawater beginning at about 13 ka during rapid post-glacial sea level rise (Yamanaka et al., 2010). At present, the outer rim of the Aira caldera circles round the northern part of Kagoshima Bay, and the main part of the caldera floor forms the seafloor at about 100 m water depth. At the southern part of the Aira caldera, Sakurajima volcano has been active since 13 ka. A geophysical investigation demonstrated that the central magma plumbing system of the Aira caldera is not located beneath Sakurajima volcano, but is beneath the Wakamiko crater (Ishihara, 1990). Fumarolic gas bubble emanations observed ubiquitously on the seafloor have been considered as evidence for the presence of the magma directly beneath the Wakamiko crater (Kikawada et al., 2007). Geochemical evidence found at several sites in the Wakamiko crater, such as anomalously high concentrations of mercury, antimony, arsenic and trace heavy metals in the surface sediment (Sakamoto, 1985), and occurrence of hydrothermal petroleum (Yamanaka et al., 1999, 2000) suggests that the hydrothermal activity has been active for a long time. The seafloor of the Wakamiko crater is covered with unconsolidated sediments of up to 80 m thick (Hayasaka, 1987), thought to be supplied from the area surrounding Kagoshima Bay. A significant amount of the terrestrial sediment is expected to be washed into Kagoshima Bay, because the Ito pyroclastic flow deposit is easily eroded (Yokoyama, 2003). The sedimentation rate within the Wakamiko crater is 0.5–4.2 mm per year (Oki, 1989). Mineralogical and geochemical studies of the Ito pyroclastic flow deposit reported that it is composed of volcanic glass (77–91%), plagioclase (2.4–19%), quartz (4.6–5.0%), and heavy minerals (2.4–4.5%) such as pyroxene, amphibole, and magnetite, and has a rhyolitic chemical composition (Oba et al., 1967). The hydrothermal activity in the Wakamiko crater is developed in a geologic setting where volcanic sediment of felsic composition is thickly deposited on a caldera floor that is semi-submerged to 200 m water depth. Four active hydrothermal sites were discovered within the Wakamiko crater during submarine dive expeditions conducted in 2003, 2005, 2007, 2008, and 2010 using the remotely operated vehicle (ROV) Hyper-Dolphin of the Japan Agency for Marine-Earth Science and Technology (JAMSTEC) (Fig. 2). Low-temperature hydrothermal activity associated with the mineralization of barite, stibnite, realgar, and pyrite at the SES site (formerly called the South site) was reported in a previous study (Ishibashi et al., 2008) (Fig. 2). Evidence for hydrothermal activity was detected as a white patch at the NSE site (formerly called the North site) (Nakaseama et al., 2008). High-temperature fluid venting at the WHV site, discovered in 2007, represents the center of the hydrothermal activity in the Wakamiko crater (Yamanaka et al., 2013). At least in years 2007, 2008, and 2010, the fluid temperature was consistently as high as 200 °C, which is close to the boiling point of seawater at this depth. During the dive expedition in 2007, another active site, the SWS site, was discovered and a widely distributed white patch associated with weak fluid shimmering was observed (Yamanaka et al., 2013).

2. Geologic setting

3. Methodology

Marine shallow-water hydrothermal activity was discovered on the seafloor of the submarine Wakamiko crater, located in the northern part of Kagoshima Bay, southwest Japan (Ishibashi et al., 2008; Nakaseama et al., 2008; Yamanaka et al., 2013; Fig. 1). The Wakamiko crater is a small depression with a parallelogram shape of 4 × 2 km at 200 m water depth (Fig. 2). The depression is the eruption center of

3.1. Sampling and processing of sediment cores Five sediment cores were collected using a piston corer during the KT-08-09 expedition of the R/V Tansei-maru in May 2008. In this study, three sediment cores were examined (Table 1, Fig. 2). The PC-1 core was collected from the vicinity of the SWS site where

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Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

Wakamiko crater 150m

Sakura -jima

50m

150m 100m

200m

250m

150m

100m 200m

50m

Aira caldera

0

10 km

Fig. 1. Bathymetric map of Kagoshima Bay (Hayasaka, 1987) showing the location of the Wakamiko crater. Dashed lines in the northern part of the Kagoshima Bay show the outer rim of the Aira caldera (Aramaki, 1984).

low-temperature fluid shimmering was observed. The PC-2 core was collected from the vicinity of the WHV site where high-temperature fluid venting (measured maximum temperature = 198.6 °C) was

discovered. The PC-5 core was collected to obtain reference sediment from the southeast corner of the crater, which is far from any known active hydrothermal sites.

Hydrothermal site WHV site NSE site

Sampling point SES site

PC-1 PC-2

SWS site

PC-5

Fig. 2. Detailed bathymetric map of the Wakamiko crater with core locations. The map is based on data from a multibeam echosounder, SeaBat 8160, during the NT05-12 and NT05-13 expeditions of the R/V Natsushima. Contours are every 5 m.

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44 Table 1 List of samples and sampling locations. Core ID

Site name

Latitude

Longitude

Depth (m)

Core length (cm)

PC-1 PC-2 PC-5

SWS site WHV site The southeast corner of the crater

31°39.51′N 31°40.07′N 31°39.27′N

130°45.89′E 130°45.70′E 130°47.52′E

200 200 200

240 340 338

Sediment sub-samples were collected at intervals of 5 to 10 cm from the cores. Sediment plugged in a core catcher was also provided for the study, since penetrations of the piston coring are often likely to cease at the layer of hydrothermally indurated sediment, which included abundant alteration minerals. The core catcher sediment was labeled as CC and its length was adjusted to 10 cm in the description, regardless of the actual recovery length. Pore fluid was squeezed from the sediment sub-samples within 24 h after recovery, similar to the procedure reported by Nakaseama et al. (2008). 3.2. Mineralogical and geochemical analyses of sediment samples Minerals in the sediment samples were identified by X-ray diffraction (XRD), Rigaku RAD II A, at the Department of Earth and Planetary Sciences, Kyushu University. The XRD measurements were conducted at 30 kV and 15 mA using Ni-filtered Cu-Kα (λ = 1.5418 Å) radiation. Step scan XRD data (2–64° 2θ, 0.05° 2θ step width, 1 s/step) were collected for bulk sediment samples. Step-scan XRD data (2–32° 2θ, 0.05° 2θ step width, 1 s/step) were collected for clay fraction (b2 μm) samples, under air-dried and ethylene glycol-saturated conditions. Depths of the samples are noted in Figs. 3, 5 and 9. Prior to separation of the clay fraction, sediment was disaggregated in distilled water and rinsed several times to remove dissolved salts. Clay

31

fractions were obtained from suspending bulk samples in distilled water for 5 h according to Stokes' law. For some samples which consist of mostly individual clay mineral, the chemical composition was determined using a transmission electron microscope (TEM) equipped with an energy dispersive spectrometer (EDS), JEOL JEM-2010FEF, in the Research Laboratory for High Voltage Electron Microscopy, Kyushu University. The TEM was operated at an accelerating voltage of 200 kV. Samples for the TEM analysis were prepared by settling on a carbon-coated copper grid after ultrasonic dispersion of powdered clay fractions in alcohol. Grains found in the PC-1 sediment were observed and analyzed using an electron probe micro-analyzer (EPMA) (JEOL JCXA-733). Determination of chemical composition of the grains was done by a wavelength dispersive spectrometer (WDS) attached to the EPMA. For this measurement, the grains were fixed with resin onto a thin section and polished. The major element composition of bulk samples was determined using X-ray fluorescence (XRF — Rigaku Supermini) at Senshu University. Prior to the measurement, the samples were dried at 750 °C and melted at 1200 °C in a mixture consisting of 0.5000 g sediment powder and 5.0000 g lithium tetraborate (Li2B4O7). The calibration lines for XRF analyses were determined based on Sato and Mizukami (2010). Accuracy of the calibration line is 0.58 wt.% and 0.26 wt.% for SiO2 and Al2O3, respectively. For Fe2O3 and CaO, accuracy of the calibration line is 0.13 wt.%. For the other elements the accuracy of the calibration line is less than 0.1 wt.%. 3.3. Measurement of oxygen and hydrogen isotopic composition of clay minerals Oxygen and hydrogen isotope values of clay minerals were determined for some clay fraction samples. We selected representative

Fig. 3. Visual description and mineral assemblage of the PC-5 core.

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Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

samples, which consisted of mostly single types of clay mineral, to obtain information on the formation temperature of the minerals. Prior to isotope measurement organic matter was removed following Jackson (1958) and free Fe-oxides were removed following Mehra and Jackson (1960). After the treatment, prepared samples were examined by XRD to confirm the purity. Isotope measurements were conducted at GNS Science. For measurement of oxygen isotope, oxygen was extracted from clay fraction samples using a CO2-laser and BrF5 following the method of Sharp (1990). Samples and standards were heated overnight to 150 °C prior to loading into the vacuum extraction line. These were then evacuated for about 6 h at room temperature. Blank BrF5 runs were done until the background yield was less than 0.2 μmol oxygen. Oxygen yields were recorded and CO2 gas analyzed on a Geo20-20 mass spectrometer. Oxygen isotope values were normalized to the international quartz standard NBS-28 using a value of + 9.6‰ or UWG-2 Garnet using a value of + 5.8‰. Oxygen values for NBS-28 or UWG-2 analyzed with the samples had values that varied by less than 0.15‰. For measurement of hydrogen isotopes, samples were pyrolyzed at 1450 °C in silver capsules. Samples were analyzed on a HEKAtech high temperature elemental analyzer coupled with a GV Instruments IsoPrime mass spectrometer. All samples were analyzed in triplicate. All results of hydrogen isotope are reported with respect to VSMOW, normalized to international standards IAEA-CH-7, NBS30 and NBS22 with reported δD values of − 100‰, − 66‰ and − 118‰. The precision on the standards was ±1.5‰ (1 sigma).

observed between 135 and 150 cmbsf in this unit, which implies that the sediment was not affected by hydrothermal activity. Unit B was characterized by pumiceous sediment associated with upward fining. Unit C was characterized by silt-sized sediment. In all units the sediment was olive black or gray in color. The PC-5 core was characterized by the abundant occurrence of volcanic glass. As shown in Fig. 4a, a large broad peak was found around 2θ = 20–30° in the XRD patterns of bulk samples, which indicates the presence of large amounts of volcanic glass. In addition to volcanic glass, the XRD analysis of the bulk samples showed the occurrence of cristobalite, plagioclase, and quartz, with small amounts of pyrite and mica clay (Fig. 4a). Similar mineral assemblages have been reported for the terrestrial sediment recognized as the Ito pyroclastic flow deposit (e.g., Oba et al., 1967). Major element compositions determined by XRF analysis of selected bulk samples are given in Table 2, with reported values for sediment collected from nearby seafloor and terrestrial areas. Four samples show high concentration of SiO2 (66–69 wt.%) and Al2O3 (14–15 wt.%) and low concentration of MgO (1.0–1.8 wt.%). This chemical composition is close to the range of felsic volcanic rock. Moreover, the chemical composition is similar to an average chemical composition of 11 sediments collected from the Kagoshima Bay (Table 2, Ref-2; Nakaseama et al., 2008) and of 11 sediments collected from the Wakamiko crater (Table 2, Ref-3; Nakaseama et al., 2008). Similar compositions have been reported for the terrestrial Ito pyroclastic flow deposit (Table 2, Ref-1; Oba et al., 1967).

3.4. Chemical analysis of pore fluids

4.2. Hydrothermally altered core from the high-temperature fluid venting site

Chemical analysis of pore fluids for some unstable species was conducted onboard within 48 h after the fluid extraction, following the protocol adopted by the ODP (Ocean Drilling Project) expeditions (Gieskes et al., 1991). Measurement of pH was conducted using a pH electrode at room temperature. Titration alkalinity was determined following the Gran method (Gieskes and Rogers, 1973). Colorimetric techniques were used for analysis of Si by the molybdenum blue method (Armstrong, 1951) and of NH4 by indophenol method (Solorzano, 1969). Other major chemical species were analyzed on-shore at Kyushu University. Concentration of Cl was determined by Ag titration following the Mohr method (Gieskes et al., 1991). Concentration of SO4 was determined after 300 times dilution by ion chromatography (Dionex, DX-100). Concentration of K was analyzed by an atomic absorption spectrophotometer (Shimadzu, AA-680) after CsCl addition and 200 times dilution of the acidified sub-samples. Concentration of other cations (Na, Mg, and Ca) were determined after 200 times dilution of the acidified sub-samples, by ICP-AES (Seiko Instruments, SPS1200AR). Analytical error for each chemical analysis was estimated from replicate analysis. The relative error was estimated to be within ±3%. Total charge balance between cations and anions for each sample showed good agreement. 4. Results 4.1. Reference core The PC-5 core was collected to obtain reference sediment from the southeast corner of the Wakamiko crater, which is far from any discovered active hydrothermal site. Fig. 3 illustrates the lithology and color of the PC-5 core with mineral assemblages determined by XRD analysis. The PC-5 core (338 cm in total length) can be classified into three units according to differences in grain size: Unit A from 0 to 220 cmbsf; Unit B from 220 to 250 cmbsf; and Unit C from 250 to 338 cmbsf. Unit A was characterized by silt- or clay-sized sediment. Upward fining from silt to clay was observed between 10 and 20 cmbsf and 30 and 40 cmbsf in Unit A. Bioturbated burrows were

The PC-2 core was collected near the WHV site where hightemperature fluid venting was discovered. The exact position of core sampling in the WHV site is unknown, but visual observation of hydrothermal minerals, such as stibnite veins at 255–270 cmbsf, confirmed that the area was hydrothermally active at depth. Moreover, immediately after bringing the core on board, the lower section (280–330 cmbsf) of the outer aluminum pipe was still warm to the touch. Even after a few hours, the interior sediment of the core catcher had a temperature of 36.5 °C. Fig. 5 shows the lithology and color of the PC-2 core, with mineral assemblages determined by XRD analysis. The PC-2 core (340 cm in total length) can be classified into four units according to sediment characteristics: Unit A from 0 to 210 cmbsf, Unit B from 210 to 270 cmbsf, Unit C from 270 to 300 cmbsf, and Unit D from 300 to 340 cmbsf. Unit A was characterized by silt- and sand-sized sediment. Sand-sized sediment was observed between 18 and 50 cmbsf associated with upward fining and between 100 and 115 cmbsf. Unit B was characterized by the occurrence of hydrothermal stibnite precipitates in sediment of silt or silty clay sizes. In sediment between 255 and 270 cmbsf, veins of stibnite crystals were present. Unit C was characterized by very indurate silt- or clay-sized sediment. Pore fluids were not squeezed from the sediments of Unit C because of the indurate sediment. Unit D was characterized by silty clay-sized sediment. The sediment of all units in the PC-2 core was mainly olive black or gray in color. The most predominant clay mineral, smectite, was recognized throughout the core except for the surface. A strong reflection at around 15 Å was found in XRD patterns of the bulk samples (Fig. 4b). The reflection expanded to around 17 Å under ethylene glycol-saturated conditions (Fig. 6a). This indicates abundant smectite in bulk samples. The 060 reflection of smectite in Units A, B, and D was found at around 1.49 Å (Fig. 7), which indicates that the smectite in these units was dioctahedral. In contrast, the 060 reflection of smectite in Unit C was found at 1.49 Å and 1.53 Å (Fig. 7), which indicates that the smectite in this unit was a mixture of dioctahedral and trioctahedral. It is also notable that the relative intensity of the 060 reflection of trioctahedral

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

a)

Pl H 2.83 2.71

M

H

H, Py

2.00

10.05

Py Py 2.43

200

1.64

Q

3.35 3.22

4.05

300

100

0

b)

20

30

40

50

14.7

Sm = smectite St = stibnite H = halite Py = pyrite

400

St

H Py

Py

Py, H

0

10

c)

20

30

40

Sm

1.63

Py

1.99 1.92

Py

2.71 2.52 2.42

St

3.13

200

St 3.58

5.64 5.05 4.48

St

Sm

2.82

H

300

100 50

60

9.61

Ke

Sm

Ke 4.58

Ke

5.64 5.07

St

0

10

3.97

St

100

St

20

Ke

Ke

1.53

St

2.54

200

3.56

St

3.15

17.0

300

Ke = kerolite St = stibnite Sm = smectite

St

3.06 2.75

400

Intensity (cps)

60

Sm

500

Intensity (cps)

10

1.50

Intensity (cps)

M = mica clay mineral C = cristobalite Q = quartz Pl = plagioclase H = halite Py = pyrite

C

30

40

50

60

d) 400

Ke = kerolite

Ke Ke

Ke

Ke

1.53

3.18

200

2.51

4.48

300

9.61

Intensity (cps)

Ke

100

0

10

20

30

40

50

60

Fig. 4. Representative XRD patterns of bulk samples. (a) 185–190 cmbsf in the PC-5 core. (b) 250–255 cmbsf in the PC-2 core. (c) Black and white colored grains from 230–240 cmbsf in the PC-1 core. (d) White colored grains from 230 to 240 cmbsf in the PC-1 core. In panels a and d, a broad peak at around 2θ = 20–30° indicates the presence of amorphous material, probably volcanic glass. Peak positions in Å.

33

smectite (at 1.53 Å) and of dioctahedral smectite (at 1.49 Å) gradually changed in relation to vertical depth. As shown in Fig. 7, the relative amount of trioctahedral smectite gradually increases as the dioctahedral smectite decreases from 270 to 300 cmbsf. The shallower section (0–55 cmbsf) of Unit A in PC-2 core predominantly contained volcaniclastic material (volcanic glass, cristobalite, and plagioclase), and no smectite was detected (Fig. 5). In contrast, below 55 cmbsf dioctahedral smectite was predominant, along with cristobalite and plagioclase. In addition, quartz, pyrite, and a very small amount of mica clay were identified, but no volcanic glass was detected. Unit B was characterized by the predominance of dioctahedral smectite and stibnite crystals. Cristobalite, plagioclase, and pyrite were also found, but much less than in Unit A and no volcanic glass was detected. Very small amounts of mica clay and kaolin mineral were identified by XRD analysis of clay fractions. Unit C was characterized by the occurrence of both dioctahedral and trioctahedral smectites. At the bottom of Unit C, dioctahedral smectite was present in small amounts and trioctahedral smectite was dominant. Moreover, mica clay (weak peaks at about 10 Å in XRD analyses) was found at the bottom of Unit C. Also pyrite was identified in Unit C, but cristobalite, plagioclase, and volcanic glass were not found. Unit D had a similar mineral assemblage to that of Unit B. It included dominant amounts of dioctahedral smectite and stibnite, some quartz, plagioclase, and pyrite, and very small amounts of mica clay, but volcanic glass was not detected. The major element composition of the smectite was determined by TEM–EDS (Table 3, Fig. 8). As shown in Fig. 8, all the smectite particles in Unit D had uniform chemical composition. Based on the Al-rich composition, the dioctahedral smectites in Unit D were montmorillonite. Most of the smectite particles in Units A and B also had uniform chemical composition, which was close to that of the montmorillonite in Unit D. Exceptionally, two data points in Unit A showed Fe-rich chemical composition which can be considered to be the composition of nontronite. A few data points in Units A and B showed a composition between nontronite and montmorillonite. Notably, most of smectite particles in Unit C had a Mg-rich chemical composition. Based on the Mg-rich composition, the trioctahedral smectites in Unit C are Mg-saponite. Some data points of Unit C had a composition between montmorillonite and saponite. This result is in agreement with the XRD result, which indicates that the smectite in Unit C is a mixture of dioctahedral and trioctahedral smectite. Major element composition of selected bulk sediment samples determined by XRF analysis is given in Table 2. Since the bulk composition of these sediments corresponds well to the elemental ratio of the clay minerals, we propose that most of the sediment was altered into one specific phase. The bulk composition of Units B and D sediments (MgO = 4.4–5.4 wt.% and Al2O3 = 20–22 wt.%) is similar to the elemental ratio of the montmorillonite determined by TEM–EDS analysis. The very high MgO composition of Unit C sediment (MgO = 20–22 wt.%) suggests that the sediment of Unit C mostly consists of saponite. 4.3. Hydrothermally altered core from the SWS site The PC-1 core was collected from the vicinity of the SWS site where low-temperature fluid shimmering was observed. The lithology and color of the core are shown in Fig. 9, with mineral assemblages determined by XRD analysis. According to differences in grain size, the PC-1 core (240 cm in total length) can be classified into two units: Unit A from 0 to 210 cmbsf and Unit B from 210 to 240 cmbsf. Unit A was characterized by silt- and sand-sized sediment. Two upward fining areas from sand to silt were observed between 65 and 95 cmbsf and 105 and 135 cmbsf in Unit A. The sediment in Unit A was mainly olive black in color. Unit B was characterized by large amounts of brecciated silt and sand of granule- to pebble-sized grains. Small sized grains (~ 5 mm) were mainly observed in the upper

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Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

Table 2 Major element composition of bulk samples determined by XRF analysis. This study Core ID

Reported chemical composition of sediment from nearby seafloor and terrestrial areas PC-5

PC-5

PC-5

PC-5

Unit

A

A

A

C

Depth (cmbsf)

25–30

100–105

212–216

290–295

Mineral

Q, Pl, C, Py

Q, M, Pl, C, Py

Pl, C

Q, M, Pl, C, Py

(a) Unaltered core Major element (wt.%) SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 Sum

65.95 0.46 13.94 5.03 0.15 1.76 2.05 5.59 1.95 0.14 97.01

68.06 0.46 14.41 4.50 0.07 1.55 1.34 4.93 1.79 0.14 97.26

67.62 0.76 15.00 4.92 0.14 1.08 3.71 4.64 2.46 0.18 100.51

68.82 0.48 15.03 4.50 0.07 1.30 1.80 4.17 2.25 0.11 98.50

Ref-1

Ref-2

Ref-3

70.7 0.25 14.2 2.58 0.08 0.57 2.23 3.31 2.60 0.07 96.4

63.9 0.6 16.4 6.2 0.2 2.1 5.3 1.7 0.18 0.0 99.9

65.0 0.6 18.3 5.8 0.1 1.5 3.1 3.1 1.8 0.22 99.9

This study Core ID

PC-2

PC-2

PC-2

PC-2

PC-2

PC-2

PC-1

Unit

A

B

C

C

D

D

B

Depth (cmbsf)

190–195

230–235

275–280

285–290

320–325

CC

CC

Mineral

Mont, Q, Pl, C, Py

Mont, Q Py, St

Sap, Mont, Pl, Py

Sap, Mont, Py

Mont, Q, Pl, Py

Mont, Q, St, Py

Ker, St

54.38 0.33 9.80 3.90 0.02 22.36 0.45 3.44 0.65 0.02 95.34

57.10 0.59 22.20 4.53 0.01 4.22 0.84 2.59 1.66 0.04 93.77

58.13 0.59 21.60 3.92 0.01 5.42 0.88 2.64 1.69 0.04 94.90

61.42 0.04 1.38 0.51 0.04 30.25 n.d. 1.24 0.20 0.01 95.07

(b) Hydrothermal altered core Major element (wt.%) 59.26 SiO2 0.58 TiO2 19.98 Al2O3 5.25 Fe2O3 MnO 0.02 MgO 4.88 CaO 1.06 4.96 Na2O 1.09 K2O 0.05 P2O5 Sum 97.13

58.98 0.59 20.98 4.16 0.01 4.44 0.65 4.54 1.40 0.03 95.78

54.53 0.36 11.85 3.54 0.01 20.00 0.42 4.10 0.80 0.02 95.64

Ref-1: average value of nearby terrestrial sediment (Oba et al., 1967). Ref-2: average value of sediment from Kagoshima bay (Nakaseama et al., 2008). Ref-3: average value of sediment from Wakamiko crater (Nakaseama et al., 2008). Total iron was expressed as Fe2O3. Q: quartz, Pl: plagioclase, C: cristobalite, Py: pyrite, M: mica, Mont: montmorillonite, Sap: saponite, St: stibnite, Ker: kerolite. n.a.: not analyzed. n.d.: not detected.

section (210–220 cmbsf) and large-sized grains (~ 10 mm) were found in the lower section (220–240 cmbsf) of Unit B. The sediment in Unit B was dark gray in color. Kerolite was predominantly found in Unit B of PC-1 core, which was characterized by strong reflection at around 9.6 Å in the XRD patterns of bulk samples (Fig. 4c, d). The reflection did not expand under ethylene glycol-saturated conditions (Fig. 6b). A 060 reflection was found at around 1.53 Å in the XRD patterns (Fig. 4c, d). TEM–EDS analyses showed that the clay minerals in Unit B have a Mg-rich chemical composition (Table 3). Since the basal spacing of talc is 9.35–9.40 Å, the clay mineral in Unit B is distinguishable from talc, and can be identified as kerolite (Brindley et al., 1977). As pointed out by Cuadros et al. (2008), “disordered, hydrated talc” should be used instead of “kerolite”, because “kerolite” is now discredited by the IMA Committee on New Minerals, Nomenclature and Classification. However, in this paper, we use “kerolite” following previous pioneering studies for clay minerals in a submarine hydrothermal system (Dekov et al., 2008a; Dekov et al., 2008b).

Mineral assemblages of the PC-1 core are given in Fig. 9. Unit A is dominated by volcaniclastic material (volcanic glass, cristobalite, and plagioclase), as well as quartz and pyrite. A large broad peak was notable at around 2θ = 20–30° in the XRD patterns of the bulk samples, which indicates the presence of abundant volcanic glass in Unit A. In the lower section (178–210 cmbsf) of Unit A, small amounts of clay minerals (smectite, mica clay and kaolin mineral) were observed by XRD analysis of the clay fraction. Sediment in Unit B was mainly composed of brecciated grains of ~ 10 mm size (Fig. 9). As shown in Fig. 4c and d, XRD analysis revealed that these grains were kerolite. As minor components, stibnite and smectite were identified, and volcanic glass was also recognized. Fine sediment (b1 mm) that constitutes the matrix of Unit B is predominantly composed of kerolite and stibnite, as identified by XRD analysis (Fig. 4c). In the fine-grained sediment of the upper section (210–220 cmbsf), smectite and quartz were present, in addition to kerolite and stibnite. Major element composition of one bulk sample of Unit B determined by XRF analysis is given in Table 2. The sample has high

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

35

Fig. 5. Visual description and mineral assemblage of the PC-2 core.

a)

dio-Sm

Sm = smectite

6000

tri-Sm

M = mica clay mineral

4000

Unit B : 250 - 255 cmbsf

5

10

Sm Sm

Sm 3.10

Sm 3.35

10.1

0

8.67

M Sm

5.19

2000

5.64

Intensity (cps)

Sm

16.8 15.5

Sm

8000

15

20

25

30

Unit C : 275 - 280 cmbsf

b) 1200

9.61

1000

Ke = kerolite

800 600

293 - 295 cmbsf

3.18

Ke

400

Ke 4.77

Intensity (cps)

285 - 290 cmbsf

Ke

200 0

5

10

15

Unit D : 320 - 325 cmbsf 20

25

30 60

Fig. 6. Representative XRD patterns of clay fraction samples. (a) 250–255 cmbsf in the PC-2 core. (b) 230–240 cmbsf in the PC-1 core. Solid line indicates XRD analyses of air-dried samples. Dashed line indicates XRD analyses of ethylene glycol saturated samples.

62

64

Fig. 7. Comparison of the 060 reflection of smectite in XRD patterns of bulk samples from the PC-2 core. A 060 reflection at 1.49 Å indicates dioctahedral smectite. A 060 reflection at 1.53 Å indicates trioctahedral smectite.

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Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

Table 3 Major element composition and structural formulae of representative clay minerals, determined by TEM-EDS analysis. Core ID

PC-2

PC-2

PC-2

PC-2

PC-2

PC-2

PC-1

Unit

(A)

(B)

(B)

(C)

(D)

(D)

(B)

Depth (cmbsf)

65–70

210–215

250–255

285–290

320–325

326–331

230–240

Mineral

Mont

Mont

Mont

Sap

Mont

Mont

Ker

65.12 n.d 24.50 3.12 5.49 0.22 0.32 0.85 0.46 100.0

59.47 n.d 8.18 3.58 26.75 0.58 0.61 0.27 0.98 100.0

62.75 n.d 26.79 3.55 5.32 0.14 0.56 0.60 0.32 100.0

60.66 0.57 25.89 1.66 5.97 0.13 1.76 1.52 2.01 100.0

66.1 n.d 2.09 2.04 29.3 n.d 0.10 0.37 0.10 100.0

3.84 0.16 4.00

3.56 0.43 4.00

3.71 0.29 4.00

3.69 0.31 4.00

3.99 0.01 4.00

1.55 0.14 0.31 2.00

0.14 0.36 2.39 2.89

1.58 0.16 0.26 2.00

1.55 0.08 0.37 2.00

0.14 0.09 2.64 2.87

0.01 0.04 0.06 0.17 0.47

0.03 0.06 0.02 0.00 0.025

0.01 0.06 0.05 0.21 0.54

0.01 0.21 0.12 0.17 0.68

0.00 0.01 0.03 0.00 0.04

(a) Major element composition of representative clay minerals Element (wt.%) 65.14 64.92 SiO2 TiO n.d 0.73 24.44 24.80 Al2O3 4.70 1.49 Fe2O3 MgO 4.39 5.89 CaO 0.20 0.20 0.36 1.50 Na2O K2O 0.40 0.41 0.47 0.30 SO3 Sum 100.0 100.2 (b) Calculated structural formulae of representative clay minerals On the basis of O10(OH)2 Tetrahedral Si 3.85 3.84 Al 0.15 0.16 Sum of tetrahedral cations 4.00 4.00 Octahedral Al 1.55 1.58 Fe 0.21 0.05 Mg 0.24 0.37 Sum of octahedral cations 2.00 2.00 Interlayer Ca 0.01 0.01 Na 0.04 0.17 K 0.03 0.03 Mg 0.15 0.15 Sum of interlayer charge 0.39 0.53

The structural formulae for pure clay minerals were calculated on the basis of 22 oxygen atoms, using the major element composition, except for TiO2 and SO3. Mont: montmorillonite, Sap: saponite, Ker: kerolite. n.d.: not detected.

concentrations of MgO (~30 wt.%) and SiO2 (~ 61 wt.%), and concentrations of K2O, CaO and TiO2 that are below limits of detection or just measureable. This result supports observations and XRD analysis that the sediment of Unit B predominantly consists of kerolite. Back-scattered electron (BSE) images of the grains in Unit B are given in Fig. 10. It is notable that stibnite is associated with the rim

of the kerolite grains (Fig. 10b). EPMA analysis revealed slight diversity in chemical composition within the grains (Table 4). While the rim had a Mg-rich chemical composition (points #1, #2, #4 in Fig. 10), the inside of a grain had a relatively Al-rich chemical composition (points #3, #5). 4.4. Pore fluid chemistry

Fe

0.8

0.2

0.4

0.6

0.4

0.6

0.2

Al

0.8

0.8

0.6

0.4

0.2

Mg

Fig. 8. Mg, Al, and Fe atomic ratios of smectite in the PC-2 core analyzed by TEM-EDS. Open diamonds represent smectite in 65–70 cmbsf in Unit A. Open circles represent smectite in 210–215 cmbsf and 250–255 cmbsf in Unit B. Solid squares represent smectite in 285–290 cmbsf in Unit C. Solid circles represent smectite in 320– 325 cmbsf and 326–331 cmbsf in Unit D.

Results of the chemical analysis of the pore fluid samples are shown as vertical profiles in Fig. 11. Pore fluids of the reference PC-5 core had a chemical composition similar to that of local seawater from the seafloor to 3 mbsf, with exception for SO4, NH4, and alkalinity (Fig. 11a). Well-correlated decrease in SO4 concentration with increase in alkalinity and NH4 concentration would reflect a sulfate reduction reaction, which is a typical organic matter digenesis in coastal sediment (e.g., Jørgensen, 1982). A previous study reported similar chemical profiles at sites within the Wakamiko crater (Nakaseama et al., 2008). Pore fluid from the PC-2 core, which was collected from the vicinity of the high-temperature fluid venting site (WHV site), showed distinctive chemistry between the upper and lower sections (Fig. 11b). Pore fluid in the upper section (Units A and B) showed basically similar chemistry to that of seawater, which suggests seawater penetration as observed in the PC-5 core. Pore fluid chemistry in the lower section (Unit D) is characterized by high Si and NH4 concentrations and low Na, Cl, Mg, and SO4 concentrations. Pore fluid could not be extracted from Unit C due to induration of the sediment. Pore fluid from the PC-1 core from the low-temperature fluid shimmering site (SWS site) also showed distinctive chemistry

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

37

Fig. 9. Visual description and mineral assemblage of the PC-1 core.

between the upper and lower sections (Fig. 11c). Pore fluid from sediments in Unit A had a similar chemistry to that of seawater, whereas pore fluids from sediments in Unit B clearly deviate from the fluid composition in Unit A. 5. Discussion 5.1. Subseafloor physical and chemical environment In order to understand the subseafloor environment where abundant Mg-clay minerals occur, we first discuss the pore fluid chemistry recovered from the cores. Pore fluid chemistry provides information reflecting geochemical processes within the sediment layer such as diagenesis of organic/inorganic matter and hydrothermal alterations. Moreover, drastic change of pore fluid chemistry would be associated with dynamic fluid flow within the sediment. As shown in Fig. 11b and c, chemical profiles of pore fluids in the vicinity of active hydrothermal fields (PC-1 and PC-2) changed at ~ 210 and ~ 270 cmbsf, respectively. On the other hand, pore fluid above these depths showed similar chemical composition to that of local seawater. These profiles are quite distinctive from the pore fluid profiles of PC-5 reference site, where well-correlated decrease in SO4 concentration with increase in alkalinity and NH4 concentration were found (Fig. 11a). Pore fluid chemistry of the lower sections (Unit D of PC-2 and Unit B of PC-1) showed higher Si and NH4 concentrations and lower Mg, Na, Cl and SO4 concentrations. Such a coordinated change in concentration of many ions is difficult to explain as a result of only fluid–sediment interaction. Rather, it could be attributed to a result of mixing with a different type of fluid. As discussed in previous studies for the surface sediment in the Wakamiko crater (Ishibashi et al., 2008; Nakaseama et al., 2008), this drastic change of pore fluid chemistry is attributed to the contribution of a hydrothermal fluid component into the deeper sediment layer. Based on geochemical data of the hydrothermal fluid venting from the WHV site (Yamanaka et al., 2013), we can estimate the chemical composition of a hypothetical mixture and compare it with the observed pore fluid chemistry. In Fig. 11b, the ion concentrations of a hypothetical fluid consisting of 1:2 mix between a hydrothermal component and seawater is plotted as arrows at the bottom of each plot. pH of the hypothetical fluid cannot be calculated due to its non-conservative property. The composition of the pore fluid in Unit D of PC-2 well agrees with that of the hypothetical fluid in

all the chemical species, with the exception of Ca and K. It is notable that low concentrations of Na and Cl of about three quarters of the composition of seawater is best explained by a contribution from the hydrothermal component, which was considered to have originated from a mixture of meteoric water and seawater (Ishibashi et al., 2008). Discrepancy in concentrations of Ca and K may be explained by ion exchange between the pore fluid and interlayer cations of the smectite. Pore fluid in Unit B of PC-1 also showed similar chemistry to that of the hypothetical mixture between the hydrothermal component and seawater. Pore fluid chemistry results revealed that sediment 2 to 3 meters below the seafloor surface in both active hydrothermal fields is likely to be occupied by pore fluid, affected by intrusion of the hydrothermal component. For PC-2, Unit C would be considered as the interface between the sediment layer occupied by pore fluid of seawater composition (Units A and B) and that affected by the hydrothermal component (Unit D), although pore fluid could not be squeezed from the sediment samples of Unit C due to sediment induration. Similarly, Unit B of PC-1 could be considered as the shallowest level of intrusion of the hydrothermal component. Pore fluid could not squeezed from the sediment samples obtained by the core catcher (230–240 cmbsf), because they dominantly consist of coarse grains. In both cases, piston core penetration unfortunately stopped at those layers. Therefore, the thickness of the layer affected by intrusion of the hydrothermal component is unknown. We have calculated the Mg/Si mole ratio of a hypothetical seawater– hydrothermal mixture, based on geochemical data presented in Yamanaka et al. (2013) (Fig. 12). Concentrations of these two elements in the reactant fluid are important to discuss the formation of Mg-rich clay minerals during fluid–sediment interaction. Seawater contains low Si and high Mg (Si = 0.16 mmol/kg and Mg = 54 mmol/kg), whereas the hydrothermal component contains high Si and zero Mg (Si = 6.4 mmol/kg). High temperature hydrothermal fluid is usually depleted in Mg, because it experiences Mg loss during the fluid circulation pathway prior to ascending to the seafloor from the fluid reservoir at depth. Therefore, the Mg/Si ratio of the pore fluid drastically decreases, as the contribution of the hydrothermal component to pore fluid increases. The Mg/Si ratio for a seawater-dominant pore fluid is above 300, 16 for the hypothetical fluid composed of a 1:2 mixture between the hydrothermal component and seawater (mixing ratio = 0.5), and less than 1 for pore fluid consist dominantly of the hydrothermal component. As shown in this calculation, intrusion of the

38

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

hydrothermal component into pore fluid significantly affects not only the temperature condition, but also the chemical environment beneath the seafloor. It is important to note that the vertical profiles of pore fluid chemistry represent only the present subseafloor environment. Since the hydrothermal fluid flows in sediment layer could be more dynamic and variable, it does not automatically mean that the identified clay minerals are formed in such an environment. However, the measured pore fluid chemistry at least endorses the possibility of a wide range of geochemical conditions within the sediment layer beneath the seafloor. Contrary to the geochemical environment, the sedimentation style is rather monotonous in the study area. The PC-5 reference core was characterized by the abundant occurrence of volcanic glass, regardless of sediment size (Fig. 3). The Wakamiko crater is located within an inner bay of an ~ 10 km in diameter and surrounded by a subaerial region where significant amount of pyroclastic flow deposits are recognized in several places. The similar chemical composition of the sediment from PC-5 core to that of the subaerial volcanic sediment (Table 2a), with a rapid sedimentation rate, suggests that the seafloor sediment dominantly consist of volcanic material once located in the subaerial region. Compared with the monotonous sediments, a steep gradient of geochemical environment is expected below the seafloor, which could be an important factor for formation (and dissolution) of clay minerals in the sediment layer.

a

b

1

5.2. Oxygen and hydrogen isotopes of hydrothermal clay minerals

2 3 c

4 5

Fig. 10. Representative images of grains found between 230 and 240 cmbsf in the PC-1 core. (a) Scan image of a thin section prepared for EPMA analysis. (b) Back-scattered electron (BSE) image by EPMA. (c) Back-scattered electron (BSE) image by EPMA. In panel a, squares indicate the area of images (b) and (c). In panels b and c, circles indicate points chemically analyzed by EPMA. The chemical composition of the indicated points is listed in Table 4. In panel b, “Stb” indicates euhedral crystals of stibnite.

Table 4 Chemical composition of grains between 230 and 240 cmbsf in the PC-1 core. Analyzed pointa

1

2

3

4

5

Element (wt.%) SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O SUM

59.3 0.03 0.12 0.13 0.00 28.3 0.12 0.18 0.03 88.2

58.5 0.04 0.43 0.58 0.00 27.8 0.11 0.21 0.04 87.7

51.6 0.10 4.14 0.54 0.02 20.4 0.41 0.43 0.14 77.7

62.49 0.02 0.58 0.47 0.07 29.35 0.13 0.12 0.06 93.28

54.17 0.19 5.05 0.75 0.00 19.37 0.10 0.67 0.16 80.46

a

Results of the measurement of oxygen and hydrogen isotope values of seven clay fraction samples are listed in Table 5. Vertical profiles of isotopic values of the smectite in the PC-2 core are shown as Fig. 13. Overall the data show that the smectite tends to have more negative δ18O and δD values with depth. Although we analyzed only one kerolite sample from the PC-1 core, the isotope value is in a similar range to the deepest smectite in the PC-2 core. Because the samples provided for the isotope measurement dominantly consist of single clay mineral, we can obtain information on formation temperature of the clay mineral type, assuming equilibrium for isotope exchange between the fluid and clay mineral. It is required that one estimates the isotope value of the fluid with which these interactions occurred to apply these geothermometers. From the discussion in the previous section, a mixture of the hydrothermal component and seawater would be a reasonable assumption. Isotopic values of the undiluted hydrothermal fluid were assumed to be δ18O = +2.5‰ and δD = −23‰, based on analyses of the hydrothermal fluid samples collected from the vent of WHV site (Yamanaka et al., 2013) and those of the local bottom seawater were reported as δ18O = +0.0‰ and δD = +2.0‰ in a previous study (Ishibashi et al., 2008). We tentatively assume that the hypothetical fluid discussed in the previous section is the fluid with which isotope equilibrium attained. Since the hypothetical fluid is considered to be 1:2 mixing between the hydrothermal component and seawater, the isotopic values are calculated to be δ18O = +0.8‰ and δD = −6.3‰. Oxygen isotope equilibration temperatures for the smectite were calculated according to the formula proposed by Sheppard and Gilg (1996): 6

Numbers of analyzed points correspond to those of circled points in Fig. 12.

−2

1000 ln α ðsmectite
waterÞ ¼ 2:55  10  T

–4:05

where T is temperature in Kelvin and α is the fractionation factor of oxygen isotopes between smectite and water. Sheppard and Gilg (1996) proposed that variations in the Al–Mg composition of the smectite will have little effect on the oxygen isotope ratios, because of the similar mass of these two elements, whereas variations in Fe content will have a greater effect. Most of the smectites in our samples showed Fe-poor composition in a narrow range (Fig. 8), therefore, this is not considered to have an effect on the isotope values.

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

39

a) PC-5 core

b) PC-2 core depth (cmbsf)

0

Unit 0

Si

Mg

Ca

Na

K

NH4

Cl

SO4

alk.

(mM)

(mM)

(mM)

(mM)

(mM)

(mM)

(mM)

(mM)

(meq)

3 30

60 8

14 360

480 6

12 0

8 450

600 0

30 0

pH 40 6

8

(A) 0-210cm

100

200 (B) 210-270cm

300

(C) 270-300cm (D) 300-340cm core catcher 330-340cm

c) PC-1 core

Fig. 11. Vertical chemical profiles of pore fluids squeezed from the sediments in the cores. (a) PC-5 core (b) PC-2 core (c) PC-1 core. Up arrow: concentration of the hypothetical fluid composed by 1:2 mixing between the hydrothermal component and seawater.

40

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

1000

Mg/Si ratio of the fluid

100

10

1

0.1

0.01 0.001

0.01

0.1

1

10

100

mixing ratio (hydrothermal component / seawater) Fig. 12. Mg/Si ratio of the hypothetical fluid formed by mixing fluids with a mixture of a hydrothermal and a seawater end-member. The X-axis shows the mixing mass ratio of the hydrothermal component to seawater. Note that both axes are shown in a logarithmic scale. Chemical composition of the hydrothermal component is represented by the hydrothermal fluid venting from the seafloor at WHV site. Data from Yamanaka et al. (2013).

Oxygen isotope equilibration temperature for the kerolite in PC-1 was calculated according to the recent experimental study of talc–seawater system (Saccocia et al., 2009), since the fractionation factor for kerolite-water has not been established. The oxygen isotope equilibrium temperature equation proposed by Saccocia et al. (2009) is: 6

−2

1000 ln αðtalc
waterÞ ¼ 11:70  10  T

3

–25:49  10  T

−1

þ 12:48:

In Table 5, calculated oxygen equilibrium temperatures for the samples using these formulas are listed. The estimated temperatures range from 118 °C for the shallowest sample to 164 °C at the deepest part of the PC-2 core. The estimated temperature for the kerolite at ~ 235 cmbsf of the PC-1 core is 211 °C. The range of the estimated temperatures broadly agrees with the subseafloor environment discussed in the previous section, which was affected by intrusion of the hydrothermal component. The observed highest temperature of the venting hydrothermal fluid was 198.6 °C and the temperature in the fluid reservoir at depth was estimated to be 230–250 °C (Yamanaka et al., 2013). Although we have arbitrarily fixed the mixing ratio between the hydrothermal component and seawater at 1:2, the estimated temperatures are not so different, even if the isotope value of the undiluted hydrothermal fluid or local seawater were applied to the calculation — the temperatures range from 110 Table 5 Oxygen and hydrogen isotope values and calculated formation temperature of representative clay minerals. Core ID

Unit

Depth (cmbsf)

Mineral

δ18O (‰)

δD (‰)

Calculated formation temperature (°C)

PC-2

(A) (A) (B) (B) (C) (D) (B)

65–70 190–195 230–235 250–255 285–290 320–325 230–240

Mont Mont Mont Mont Sap Mont Ker

13.5 12.1 12.5 12.4 10.2 10.3 10.6

−54.8 −61.2 −63.1 −71.6 −72.6 −69.4 −68.2

118a 135a 130a 132a 164a 163a 211b

PC-1

Mont: montmorillonite, Sap: saponite, Ker: kerolite. a Calculated following Sheppard and Gilg (1996). b Calculated following Saccocia et al. (2009).

to 140 °C for the shallowest sample and from 150 to 190 °C for the deepest sample of the PC-2 core, and from 203 to 230 °C for the deepest sample of the PC-1 core. To simplify the following discussion, we assume that the temperatures calculated in Table 5 reflect the temperatures of formation, accepting that there is some uncertainty. Hydrogen isotope equilibration temperatures can be calculated independently in a similar way using fractionation factors between the fluid and smectite proposed by Yeh (1980) and Capuano (1992). However, the calculated temperatures are unreasonably low ranging from 40 °C to negative temperatures. Moreover, according to the formula by Yeh (1980) and Capuano (1992), the calculated temperatures show a profile where higher temperatures occur toward the seafloor. Marumo et al. (1995) reported the hydrogen isotope values of smectite collected from an active geothermal field in Noboribetsu in Japan. They demonstrated that fractionation of hydrogen isotope value between smectite in the geothermal field (δD = −71 to −65‰) and hot spring waters of meteoric origin (from −50 to −40‰) was in isotope equilibrium at temperatures considered possible for smectite formation (from 100 to 150 °C). Uysal et al. (2000) studied authigenic clay minerals, which consisted predominantly of the mixed layer illite/smectite, collected from coal measures in Bowen Basin in Australia. They reported that the hydrogen isotope values of the clay minerals (δD = −110 to −77‰) were likely to be too low (depleted in δD), because estimates of isotopic values of the reactant fluid using the palaeotemperatures based on virtrinite reflectance were in disagreement with the range expected from a deeply buried sediment basin. They mentioned the possibility that post-formation isotope exchange may have occurred. Within the Wakamiko crater, occurrence of hydrothermal petroleum is reported and considered a product of hydrothermal maturation of sedimentary organic matter (Yamanaka et al., 1999, 2000). In the case of the samples in this study, it is possible that the hydrogen isotope values were affected by a similar post-formation isotope exchange. 5.3. Montmorillonite formation process As summarized in Fig. 5, the PC-2 core collected from the vicinity of the high-temperature venting site is characterized by the abundant occurrence of smectite. Abundant smectite was not observed in the reference core (PC-5). Based on the estimated formation temperature proposed in the previous section (118–163 °C), the abundant montmorillonite in Units B and D of the PC-2 core is attributed to be formed by hydrothermal process. According to the previous mineralogical study on subaerial sediment around the Kagoshima Bay (Miyauchi et al., 1972), montmorillonite found in subaerial sediment was identified by very weak 001 reflection in an XRD scan. On the other hand, the montmorillonites in the PC-2 core showed strong 001 reflection in XRD scans of bulk samples (Fig. 4b). This suggests that the montmorillonite in the PC-2 core did not originate from subaerial sediment, but formed in situ after sedimentation of the original sediment. Occurrence of notable amount of stibnite (Sb2S3) as veins in Units B and D is evidence that hydrothermal processes have occurred within the sediment layer (Fig. 4b). Stibnite precipitation results from a decrease in fluid temperature, based on experimental studies (e.g., Krupp, 1988) and on observations of a natural system (e.g., Neiva et al., 2008). Neiva et al. (2008) demonstrated that stibnite precipitated by fluid cooling from 225 to 128 °C in antimony quartz veins, based on analysis of fluid inclusions in the quartz. The estimated temperature of formation of the montmorillonite is broadly in agreement with the temperature range where stibnite preferably precipitates. This accordance suggests a hydrothermal origin of the montmorillonite related to ascension of hydrothermal fluid to the 55 cmbsf layer. Taking account of the abundant occurrence of volcanic glass in the reference core (PC-5), it seems reasonable to assume that the

Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

a) δ18O values

41

b) δD values (‰, VSMOW)

5 (cmbsf)

10

15

(‰, VSMOW)

20

0

-80 (cmbsf)

-70

-60

-50

0

50

50

100

100

150

150

200

200

250

250

300

300

350

350

Fig. 13. Vertical profiles of isotope composition of clay minerals in the cores. (a) δ18O values (‰, VSMOW). (b) δD values (‰, VSMOW).

montmorillonite is formed from altered volcanic glass. It is well known that montmorillonite is formed by relatively low temperatures of hydrothermal alteration from volcanic glass. Volcanic glass originated from the Aira caldera is known to have a low Mg content (MgO = 0.12 wt.%; Aoki and Machida, 2006). On the other hand, bulk chemical composition of the altered layer (Units B and D) in the PC-2 core showed substantially higher Mg content (about 5 wt.% as MgO; Table 3) than that of unaltered sediment (b 2 wt.% as MgO; Table 2a). This enrichment would be attributed to uptake of Mg from seawater during montmorillonite formation. Hocking et al. (2010) demonstrated a similar comparison of chemical composition of hydrothermal clay minerals that dominantly consisted of montmorillonite, for a shallow marine hydrothermal system in the Calypso Hydrothermal Vent Field in New Zealand. They reported higher Mg content in the submarine hydrothermal system than those in the nearby subaerial geothermal field in the Taupo Volcanic Zone, and suggested that this difference reflects substantially higher Mg concentration of seawater than that of meteoric water. Therefore, the montmorillonite in the PC-2 is attributed to be formed by hydrothermal alteration of volcanic glass involving Mg uptake from seawater. 5.4. Saponite formation process As shown in Fig. 5, Unit C of the PC-2 core was dominated by saponite with minor montmorillonite. Formation temperature estimated from the oxygen isotopic values is 164 °C, which suggests hydrothermal origin of the saponite as well as the montmorillonite. In previous studies on clay minerals found in submarine hydrothermal fields, saponite is in many cases attributed to be formed by hydrothermal alteration of basalt rocks (e.g., Alt et al., 1983; Shau and Peacor, 1992; Turner et al., 1993; Percival and Ames, 1993; Goodfellow et al., 1993; Giorgetti et al., 2001; Buatier et al., 2002; Giorgetti et al., 2009). Seyfried and Bischoff (1979) demonstrated that Fe-rich saponite (8.7–10.9 wt.% as Fe2O3) was the dominant alteration product of basalt glass (Fe2O3 = 13.29 wt.%) during their experimental study conducted at 150 °C. Shau and Peacor (1992) reported dominant occurrence of Fe-rich saponite (Fe2O3 = 12.4–14.1 wt.%) in the transition zone underlying the pillow lava layer (572–624 mbsf) where formation temperature of the saponite was estimated to be 60–110 °C. In comparison with these saponites altered from basalt rocks, the saponite in the PC-2 core showed substantially low content of Fe (1.9–5.4 wt.% as Fe2O3; Table 3). The saponite in the PC-2 core is unlikely to be formed by hydrothermal alteration of basalt rocks,

considering the geologic setting of the Kagoshima Bay where exposure of basaltic lava is rare. Occurrence of saponite was identified specifically within the 30 cm thick layer of Unit C in the PC-2 core. One possible formation process for the saponite is hydrothermal alteration from volcanic glass, under favorable conditions. Giorgetti et al. (2006) identified saponite as well as montmorillonite in cracks and vesicles of hydrothermally altered dacitic lava. Based on SEM and TEM observations, they recognized that saponite and montmorillonite have common textures and interpreted that both have formed by similar process of dissolution and crystallization from the volcanic glass. Monecke et al. (2007) proposed that availability of Mg is a control factor to distinguish formation of saponite and montmorillonite; saponite formation during glass dissolution should be associated with involvement of more Mg influx. If it is the case, saponite formation would be related to more seawater-dominant condition than montmorillonite formation. An alternative idea is that saponite formed by transformation of material of different composition that had been deposited in this layer. For the formation of Mg-saponite, a high content of Mg in the precursor would be preferable. As discussed in the previous paragraph, basaltic glass is unlikely to be a precursor in the study field. Kerolite may be such a precursor, because occurrence of kerolite was confirmed in another site. However, the kerolite layer found in Unit B of the PC-1 core consists of coarse grains. It may be difficult to explain that the indurated sediment layer formed from such coarse grain sediment. Abundant montmorillonite could be another candidate for such a precursor. Some previous experimental studies reported that formation of saponite was associated with dissolution of montmorillonite in pure water at hydrothermal conditions (e.g., Beaufort et al., 2001). However, it is well known that montmorillonite is relatively stable in seawater at ambient temperatures, but that transformation of montmorillonite to saponite could be possible when there is substantial change in conditions. For example, temperatures higher than formation temperatures of montmorillonite could contribute to enhanced dissolution of montmorillonite, which are present in substantial amounts in adjacent layers (Units B and D). The analytical results of the PC-2 core sediment are supportive of the latter model. Oxygen isotopic values of smectite in Units C and D were 18 O-depleted compared with those in Unit B, which corresponds to higher estimated formation temperature (Table 5, Fig. 13). Moreover, gradual change of the relative amount of saponite and montmorillonite with depth within Unit C (Fig. 7) is in accordance with the transformation process. Gradual increase of saponite in the deeper part is not in accordance with the profiles of pore fluid chemistry which showed high

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Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

Mg/Si in Unit B and low Mg/Si in Unit D. Therefore, we propose that the saponite in PC-2 core is attributed to transformation of pre-existing montmorillonite. 5.5. Kerolite formation process As summarized in Fig. 9, the PC-1 core collected from the low-temperature fluid shimmering site is characterized by the abundant occurrence of kerolite in Unit B from 210 to 240 cmbsf. Formation temperature for kerolite, based on an oxygen isotope value, is estimated to be around 211 °C (Table 5). In previous studies, formation of kerolite (or Mg-talc) is usually attributed to precipitation due to mixing of Si-rich hydrothermal fluid and Mg-rich seawater (Zierenberg and Shanks, 1994; Dekov et al., 2008a). The chemical composition of Unit B is in accordance with the idea that abundant kerolite formed by hydrothermal precipitation. XRF analysis indicated high concentrations of Mg (30 wt.% as MgO) and Si (61 wt.% as SiO2), but low concentrations of K, Ca and Ti (Table 2b). In particular, there is a notable depletion in Ti, which would be evidence for precipitation from the reactant fluid, because this element is usually immobile during hydrothermal alteration. Occurrence of kerolite as coarse grains of 1–5 mm (rarely up to 10 mm) in size was reported for the Grimsey shallow-water hydrothermal field in sediment covered spreading centers north of Iceland (Dekov et al., 2008a). Dekov et al. (2008a) proposed that the kerolite grains were buried in the sediment by collapse of hydrothermal chimneys including kerolite that had been formed as precipitates on the seafloor. In their study, formation temperatures of the kerolite was estimated to be 300–330 °C, based on oxygen isotope values and using the fractionation equation of Saccocia et al. (2009). Within the Wakamiko crater, a few high-temperature vents associated with a small chimney structure were discovered at the WHV site, which is located about 2 km north of the SWS site where the PC-1 core was collected (Yamanaka et al., 2013). At this site, hydrothermal agglomerates around the vents were composed of abundant kerolite and carbonate with minor anhydrite. Nonetheless, we propose in situ formation of the kerolite within the sediment layer, rather than sedimentation of the collapsed material transported from such a far region. The XRD analysis of powdered grains found in this unit indicates that a small amount of smectite and unaltered volcanic glass remained (Fig. 4c and d). EPMA analysis of the grains demonstrates co-precipitation of kerolite and stibnite on the rim of a grain (Fig. 10). Similar occurrence of talc associated with smectite-like phyllosilicate was documented for altered sediment collected from a hydrothermal field in Escanaba Trough (Zierenberg and Shanks, 1994).

clays at different depths would indicate that Mg-rich chlorite was formed at higher temperatures than Mg-rich smectite, which is supported by the estimated formation temperature of higher than 250 °C, based on the oxygen isotope values of the chlorite in the ODP drilling cores (Lackschewitz, 2000). In contrast, relatively few studies have reported the occurrence of Mg-rich clays in submarine hydrothermal fields related to felsic magma activity in arc–backarc settings. Marumo and Hattori (1999) documented occurrence of Mg-rich chlorite in a limited area in the JADE hydrothermal field in the Okinawa Trough, while Al-rich clay minerals, kaolinite or halloysite, were dominantly found in a wide area. Giorgetti et al. (2006) investigated the first example of occurrence of saponite in such geologic setting in the PACMANUS hydrothermal field in the Manus Basin. They found occurrence of saponite as well as montmorillonite in cracks and vesicles of hydrothermally altered dacitic lava and interpreted that both formed by similar processes of dissolution and crystallization from the volcanic glass. Our study has revealed that the Mg-rich clay minerals, saponite and kerolite, occur in felsic volcanic sediments within the active hydrothermal fields of the Wakamiko crater. As discussed in previous sections, seawater acts as an important Mg source during the hydrothermal alterations in such geologic settings. Actually, very few mineralogical studies of subaerial geothermal fields related to felsic magma activity have reported the occurrence of Mg-rich clay minerals, such as chlorite and chlorite–smectite mixed-layer. They are typically dominated by K- and Al-rich clay minerals, such as illite and illite–smectite mixed-layer (e.g., Inoue et al., 1999; Inoue et al., 2004). The vertical profile of pore fluid chemistry shown in this study suggests that occurrence of hydrothermal Mg-rich clay minerals would be restricted to the interface between sediment occupied with pore fluids of the hydrothermal component and of seawater composition, probably due to availability of Mg in seawater. Fig. 14 shows a ternary plot of Si–Mg–Al atomic ratios of the hydrothermal clay minerals analyzed by TEM–EDM, which will aid this discussion. In Fig. 14 data of volcanic glass in AT (Aira-Tanzawa) tephra that erupted during the main eruption of the Aira caldera formation is also plotted to represent precursor material prior to hydrothermal alteration (data from Aoki and Machida, 2006). For the PC-1 core, occurrence of kerolite was limited in Unit B which is considered to be the shallowest level of intrusion of the hydrothermal component. The measured Mg/Si ratio of the kerolite was

Al

5.6. Mg-rich clay mineral formation in seafloor hydrothermal systems Occurrence of Mg-rich clay minerals has been found ubiquitously in sediment-rich hydrothermal systems in mid-ocean ridge settings. For example, Mg-rich chlorite, Mg-rich chlorite–smectite mixed-layer mineral, Mg-rich talc and Mg-rich smectite occurred in sediment at Escanaba Trough and Middle Valley (Turner et al., 1993; Zierenberg and Shanks, 1994; Lackschewitz, 2000). These Mg-rich clay minerals were considered to be formed by mixing of Mg-rich seawater with Mg-depleted and Si-rich hydrothermal fluid. Zierenberg and Shanks (1994) reported that Mg-rich chlorite was formed by replacement of detrital minerals, while Mg-rich talc was formed by direct precipitation, due to mixing of seawater with hydrothermal fluid at Escanaba Trough. Zierenberg and Shanks (1994) suggest that higher silica activity favors the formation of talc relative to Mg-rich smectite. In the Middle Valley, Mg-rich chlorite or chlorite–smectite mixed-layer minerals were found at deep depths by ODP drilling (Lackschewitz, 2000), while Mg-rich smectite was found in surface sediments of push-core samples (b 30 cm depth) (Turner et al., 1993). Occurrence of the two Mg-rich

0.2

0.8

0.6

0.4

0.4

0.6

montmorillonite

0.8

0.2

volcanic glass

saponite Mg

Si 0.2

0.4

kerolite

0.8

Fig. 14. Ternary diagram showing Si, Mg and Al atomic ratios of the clay minerals analyzed by TEM-EDS. Data of volcanic glass in AT tephra is also plotted to represent precursor material. Data from Aoki and Machida (2006).

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about 1.5. If kerolite is precipitated from the pore fluid with this Mg/Si ratio, the pore fluid should be dominated by the hydrothermal component (the hydrothermal component to seawater is about 5 in Fig. 12). As discussed in previous studies, such a mixture would induce supersaturation for kerolite. A relatively high estimated temperature of formation for kerolite (211 °C) would support this formation model. Although the observed pore fluid in Unit B did not show chemistry of a hydrothermal dominant component, it is probably due to difference in the sampling depth. Pore fluid could, however, not be extracted from the sediment that dominantly consist of kerolite (coarse grains), and the pore fluid data shown in Fig. 11c came from sediment 10 cm above the kerolite-dominant sediment. Formation of kerolite by precipitation could be attributed to hydrothermal interaction at rather high water-rock ratio conditions where Al supply from silicate minerals is insufficient. Unit B sediment that consisted mainly of coarse grains would be favorable for such hydrothermal interactions, because it could provide space for rapid influx of the hydrothermal component, which causes kerolite precipitation and possibly dissolution of the primary minerals and volcanic glass. Similar coarse grain sediments were found in the reference site (220–250 cmbsf of the PC-5 core) and may be related to turbidite sedimentation from the subaerial area. For the PC-2 core, occurrence of saponite was limited in Unit C, which is considered to be the interface layer between the sediment layer occupied by pore fluid of seawater composition (Units A and B) and that was affected by the hydrothermal component intrusion (Unit D). The saponite in this Unit is a Mg-saponite, which shows high Mg/Si and Mg/Al ratios. Although the Mg/Si ratio of the saponite is close to that of the kerolite in PC-1 core, a substantially higher Mg/Si ratio of the reactant fluid is required for saponite formation. Water-rock ratio for saponite formation should be low (rock-dominant condition), because dissolved Al in the reactant fluid is very limited (micromolar level even in the hydrothermal fluid). Therefore, the availability of Mg from the reactant fluid is a critical factor for formation of saponite. Cuadros et al. (2013) reported saponite formation from the alteration of volcanic glass during an experimental study at room temperature. They confirmed the occurrence of saponite during the interaction with hypersaline fluid that contains 13.3 times more Mg than seawater, but observed only montmorillonite during interaction with seawater. Transformation of montmorillonite to saponite must be related to substantial change in conditions. High temperature is the most plausible factor, however, it is not consistent with a high Mg availability in the geochemical environment controlled by mixing of the hydrothermal component and seawater. As shown in Fig. 12, Mg/Si ratio decreases drastically from the seawater-dominant condition (implies low temperature) to the hydrothermal-component dominant condition (implies high temperature). Therefore, an adequate temperature and geochemical environment for the transformation that results in saponite formation would be limited. Occurrence of saponite only in the interface layer is a possible consequence of the specific conditions required to form saponite. Verification of the required conditions of saponite formation is not possible due to the absence of experimental data. pH is another factor that may be an important control on the stability of saponite. The pore fluid in Unit D, which is dominated by the hydrothermal component, had a lower pH than ambient pore fluid. Low pH of the reactant fluid would enhance dissolution of montmorillonite, however, it would be unfavorable for saponite formation. This means that the relatively low pH of the hydrothermal component creates a smaller window that is a conductive geochemical environment for transformation of montmorillonite to saponite. In reality, the upward flow of the hydrothermal component and the downward penetration of seawater would be a dynamic system. Within the interface layer, the fluid, temperature and mixing ratio would fluctuate. Temperature can also fluctuate, by conductive heating or cooling, which is not perfectly linked to fluid flow. In contrast, the mineral surface is relatively static. If ions released by

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dissolution of a mineral precursor are not diffused quickly, but stacked around the mineral surface, they could be available for formation of a different mineral when the temperature shifts. Cuadros et al. (2013) demonstrated that microbe activity influenced the formation of smectite from volcanic glass and that biofilms of low permeability acted as a barrier to contain different water chemistry around the glass. In the sediment layer mobility of dissolved ions would be restricted, which may enhance transformation of clay minerals in response to fluctuation in temperature and geochemical conditions. Both kerolite and saponite, and also montmorillonite are attributed to be formed by hydrothermal interactions where significant amount of seawater is involved as a Mg source. We propose that the occurrence of Mg-rich clay minerals would be strong evidence for significant seawater penetration as well as intrusion of the hydrothermal component into sediment layer of felsic composition. Occurrence of Mg-rich clay minerals provides important clues to the hydrological structure in sedimentcovered hydrothermal systems in an arc volcanic caldera setting. 6. Conclusion A mineralogical study of sediment samples revealed the occurrence of Mg-rich clay minerals, saponite and kerolite, in a volcanic sediment layer of felsic composition within the active hydrothermal fields in the Wakamiko crater. Formation temperature estimated from the oxygen isotopic values of these clay minerals indicated that both saponite and kerolite are formed by hydrothermal interactions, as well as montmorillonite that is another dominant clay minerals in this area. Pore fluid chemistry suggests that occurrence of both saponite and kerolite are related to the interface between sediment occupied with pore fluids that have a hydrothermal and seawater component. Seawater is an important Mg source for hydrothermal interactions. Comparison of Mg:Si between the clay minerals and pore fluid constrain the formation processes. Formation of saponite is attributed to hydrothermal interaction at low water-rock conditions between seawater-dominant pore fluid and the montmorillonite which had been formed at a prior stage. Kerolite is formed by precipitation caused by mixing the hydrothermal component and seawater at high water-rock conditions. Occurrence of Mg-rich clay minerals provides important clues to the hydrological structure in sediment-covered hydrothermal systems in an arc volcanic caldera setting. Acknowledgments We are grateful to the Captain and crew of R/V Tansei-maru for their skilful operations. We thank the Research Laboratory for High Voltage Electron Microscopy at the Kyushu University for TEM analysis and Andy Phillips and Jannine Cooper at GNS Science for their assistance with stable isotope analyses. We thank Dr. Javier Cuadros and two anonymous reviewers for their constructive criticism and comments on an earlier version of the manuscript. This article is part of a doctoral thesis of the first author (Y.M.). Expenses of the first author to join the research cruise were supported by the Matsumoto Scholarship Fund of Kyushu University. This study was partly supported by the “TAIGA project” that was funded by the Grant-in-Aid for Scientific Research on Innovative Areas of MEXT, and also by KAKENHI #24340135 of JSPS, Japan. References Alt, J.C., Laverne, C., Muehlenbachs, K., 1983. Alteration of the upper oceanic crust: mineralogy and processes in deep sea drilling project Hole 504B, Leg 83. Initial Reports of the Deep Sea Drilling Project 83, 217–247. Aoki, K., Machida, H., 2006. Major element composition of volcanic glass shards in the late Quaternary widespread tephras in Japan — distinction of tephras using K2O– TiO2 diagrams. Bulletin. Geological Survey of Japan 57 (7/8), 239–258 (in Japanese with English abstract). Aramaki, S., 1984. Formation of the Aira caldera, Southern Kyushu, approximately 22,000 years ago. Journal of Geophysical Research 89, 8485–8501.

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Y. Miyoshi et al. / Chemical Geology 355 (2013) 28–44

Armstrong, 1951. The determination of silicate in sea water. Journal of the Marine Biological Association of the United Kingdom 30, 149–160. Beaufort, T., Berger, G., Lacharpagne, L.C., Meunier, A., 2001. An experimental alteration of montmorillonite to a di + trioctahedral smectite assemblage at 100 and 200 °C. Clay Minerals 36, 211–225. Brindley, G.W., Bish, D.L., Wan, H.M., 1977. The nature of kerolite, its relation to talc and stevensite. Mineralogical Magazine 41, 443–452. Buatier, M.D., Fruh-Green, G.L., Karpoff, A.M., 1995. Mechanisms of Mg-phyllosilicate formation in a hydrothermal system at a sedimented ridge (Middle Valley, Juan de Fuca). Contributions to Mineralogy and Petrology 122, 134–151. Buatier, M.D., Karpoff, A.M., Charpentier, D., 2002. Clays and zeolite authigenesis in sediments from the flank of the Juan de Fuca Ridge. Clay Minerals 37, 143–155. Capuano, R.M., 1992. The temperature dependence of H-isotope fractionation between clay minerals and water: evidence from a geopressured systems. Geochimica et Cosmochimica Acta 56, 2547–2554. Cuadros, J., Dekov, V.M., Fiore, S., 2008. Crystal chemistry of the mixed-layer sequence talc–talc–smectite–smectite from submarine hydrothermal vents. American Mineralogist 93, 1228–1348. Cuadros, J., Afsin, B., Jadubansa, P., Ardakani, M., Ascaso, C., Wierzchos, J., 2013. Microbial and inorganic control on the composition of clay from volcanic glass alteration experiments. American Mineralogist 98, 319–334. Dekov, V., Scholten, J., Garbe-Schonberg, C.-D., Botz, R., Cuadros, J., Schmidt, M., Stoffers, P., 2008a. Hydrothermal sediment alteration at a seafloor vent field: Grimsey Graben, Tjornes Fracture Zone, north of Iceland. Journal of Geophysical Research 113, B11101. Dekov, V.M., Cuadros, J., Shanks, W.C., Koski, R.A., 2008b. Deposition of talc, kerolite– smectite, smectite at seafloor hydrothermal vent fields: evidence from mineralogical, geochemical and oxygen isotope studies. Chemical Geology 247, 171–194. Gieskes, J.M., Rogers, W.C., 1973. Alkalinity determination in interstitial waters of marine sediments. Journal of Sedimentary Petrology 43, 272–277. Gieskes, J.M., Gamo, T., Brumsack, H., 1991. Chemical methods for interstitial water analysis aboard Joides Resolution. Ocean Drilling Program Technical Note, 15. Giorgetti, G., Marescotti, P., Cabella, R., Lucchetti, G., 2001. Clay mineral mixtures as alteration products in pillow basalts from the eastern flank of Juan de Fuca Ridge: a TEM–AEM study. Clay Minerals 36, 75–91. Giorgetti, G., Monecke, T., Kleeberg, R., Hannington, M.D., 2006. Low-temperature hydrothermal alteration of silicic glass at the PACMANUS hydrothermal vent field, Manus Basin: an XRD, SEM and AEM–TEM study. Clays and Clay Minerals 54 (2), 240–251. Giorgetti, G., Monecke, T., Kleeberg, R., Hannington, M.D., 2009. Low temperature hydrothermal alteration of trachybasalt at conical seamount, Papua New Guinea: formation of smectite and metastable precursor phases. Clays and Clay Minerals 57 (6), 725–741. Goodfellow, W.D., Grapes, K., Cameron, B., Franklin, J.M., 1993. Hydrothermal alteration associated with massive sulfide deposits, Middle Valley, northern Juan de Fuca Ridge. Canadian Mineralogist 31, 1025–1060. Hayasaka, S., 1987. Geologic structure of Kagoshima Bay, south Kyushu, Japan. Monograph of Association for the Geological Collaboration in Japan 33, 225–233 (in Japanese with English abstract). Hocking, M.W.A., Hannington, M.D., Percival, J.B., Stoffers, P., Schwarz-Schampera, U., de Ronde, C.E.J., 2010. Clay alteration of volcaniclastic material in a submarine geothermal system, Bay of Plenty, New Zealand. Journal of Volcanology and Geothermal Research 191, 180–192. Inoue, A., Utada, M., Shmizu, M., 1999. Mineral–fluid interactions in the Sumikawa geothermal system, northeast Japan. Resource Geology Special Issue 20, 79–98. Inoue, A., Meunier, A., Beaufort, D., 2004. Illite–smectite mixed-layer minerals in felsic volcaniclastic rocks from drill cores, Kakkonda, Japan. Clays and Clay Minerals 52 (1), 66–84. Ishibashi, J.-I., Nakaseama, M., Seguchi, M., Yamashita, T., Doi, S., Sakamoto, T., Shimada, K., Shimada, N., Noguchi, T., Oomori, T., Kusakabe, M., Yamanaka, T., 2008. Marine shallow-water hydrothermal activity and mineralization at the Wakamiko crater in Kagoshima Bay, south Kyushu, Japan. Journal of Volcanology and Geothermal Research 173, 84–98. Ishihara, K., 1990. Magma transport and storage. In: Ryan, M.P. (Ed.), Pressure Sources and Induced Ground Deformation Associated with Explosive Eruptions at an Andesitic Volcano: Sakurajima Volcano, Japan. John Wiley and Sons, Chichester, New York, Brisbane, Toronto, Singapore, pp. 335–360. Jackson, M.L., 1958. Soil Chemical Analysis.Prentice-Hall, INC., Englewood Cliffs, N. J. Jørgensen, B.B., 1982. Mineralization of organic matter in the sea bed — the role of sulphate reduction. Nature 296 (5858), 643–645. Kikawada, Y., Osska, J., Oi, T., Hashimoto, J., 2007. Chemical compositions of volcanic gases from submarine fumaroles in Wakamiko Caldera, northern Kagoshima Bay, Japan. Journal of the Balneological Society of Japan 57, 30–41 (in Japanese with English abstract). Krupp, R.E., 1988. Solubility of stibnite in hydrogen-sulfide solutions, speciation, and equilibrium-constants, from 25 to 350 °C. Geochimica et Cosmochimica Acta 52, 3005–3015. Lackschewitz, K.S., 2000. Formation and transformation of clay minerals in the hydrothermal deposits of Middle Valley, Juan de Fuca Ridge, ODP Leg 169. Economic Geology 95, 361–390. Marumo, K., Hattori, K., 1999. Seafloor hydrothermal clay alteration at Jade in the backarc Okinawa Trough: mineralogy, geochemistry and isotope characteristics. Geochimica et Cosmochimica Acta 63 (18), 2785–2804.

Marumo, K., Longstaffe, F.J., Matsubaya, O., 1995. Stable isotope geochemistry of clay minerals from fossil and active hydrothermal systems, southwestern Hokkaido, Japan. Geochimica et Cosmochimica Acta 59 (12), 2545–2559. Mehra, O.P., Jackson, M.L., 1960. Iron oxide removal from soils and clays by a dithionite–citrate system buffered with sodium bicarbonate. Clays and Clay Minerals 7, 317–327. Miyauchi, N., Ogi, S., Haruyama, M., 1972. Clay minerals in the so-called “Shirasu” bed at several places. Bulletin of the Faculty of Agriculture, 22. Kagoshima University, pp. 123–130 (in Japanese). Monecke, T., Giorgetti, G., Scholtysek, O., Keeberg, R., Götze, J., Hannington, M.D., Petersen, S., 2007. Textural and mineralogical changes associated with the incipient hydrothermal alteration of glassy dacite at the submarine PACMANUS hydrothermal system, eastern Manus Basin. Journal of Volcanology and Geothermal Research 160, 23–41. Nakaseama, M., Ishibashi, J.-I., Ogawa, K., Hamasaki, H., Fujino, K., Yamanaka, T., 2008. Fluid–sediment interaction in a marine shallow-water hydrothermal system in the Wakamiko submarine crater, south Kyushu, Japan. Resource Geology 58 (3), 289–300. Neiva, A.M.R., Andras, P., Ramos, J.M.F., 2008. Antimony quartz and antimony–gold quartz veins from northern Portugal. Ore Geology Reviews 34, 533–546. Oba, N., Tsuyuki, T., Ebihara, H., 1967. Mineral and chemical compositions, and genesis of the Shirasu (I). Journal of Mineralogy, Petrology, and Economic Geology 58 (3), 81–97 (in Japanese with English abstract). Oki, K., 1989. Ecological analysis of benthonic foraminifera in Kagoshima Bay, South Kyushu, Japan. South Pacific Study 10, 1–191. Percival, J.B., Ames, D.E., 1993. Clay mineralogy of active hydrothermal chimneys and an associated mound, Middle Valley, Northern Juan de Fuca Ridge. Canadian Mineralogist 31, 957–971. Saccocia, P.J., Seewald, J.S., Shanks, W.C., 2009. Oxygen and hydrogen isotope fractionation in serpentine–water and talc–water systems from 250 to 450 °C, 50 MPa. Geochimica et Cosmochimica Acta 73, 6789–6804. Sakamoto, H., 1985. The distribution of mercury, arsenic, and antimony in sediments of Kagoshima Bay. Bulletin of the Chemical Society of Japan 58, 580–587. Sato, H., Mizukami, K., 2010. Calibration lines for quantitative analyses of major elements in ceramics samples with X-rayfluorescence analyzer. Bulletin of the Institute of Natural Sciences, Senshu University No. 41, 25–30 (in Japanese). Seyfried, W.E., Bischoff, J.L., 1979. Low temperature basalt alteration by seawater: experimental study at 70 °C and 150 °C. Geochimica et Cosmochimica Acta 43, 1937–1947. Sharp, Z.D., 1990. Laser-based microanalytical method for the in situ determination of oxygen isotope ratios of silicates and oxides. Geochimica et Cosmochimica Acta 54, 1353–1357. Shau, Y.-H., Peacor, D.R., 1992. Phyllosilicates in hydrothermally altered basalts from DSDP Hole 504B. Leg 83 - a TEM and AEM study, Contrib. Mineral. Petrol., Vol. 12, pp. 119–133. Sheppard, S.M.F., Gilg, H.A., 1996. Stable isotope geochemistry of clay minerals. Clay Minerals 31, 1–24. Solorzano, L., 1969. Determination of ammonia in natural waters by phenolhypochlorite method. Limnology and Oceanography 14, 799–801. Turner, R.J.W., Ames, D.E., Franklin, J.M., Goodfellow, W.D., Leitch, C.H.B., Hoy, T., 1993. Character of active hydrothermal mounds and nearby altered hemipelagic sediments in the hydrothermal areas of Middle Valley, northern Juan de Fuca Ridge: Data on shallow cores. Canadian Mineralogist 31, 973–995. Uysal, I.T., Golding, S.D., Baublys, K., 2000. Stable isotope geochemistry of authigenic clay minerals from Late Permian coal measures, Queensland, Australia: implications for the evolution of the Bowen Basin. Earth and Planetary Science Letters 180, 149–162. Yamanaka, T., Mizota, C., Murae, T., Hashimoto, J., 1999. A currently forming petroleum associated with hydrothermal mineralization in a submarine caldera, Kagoshima Bay, Japan. Geochemical Journal 33, 335–367. Yamanaka, T., Ishibashi, J., Hashimoto, J., 2000. Organic geochemistry of hydrothermal petroleum generated in the submarine Wakamiko caldera, southern Kyushu, Japan. Organic Geochemistry 31, 1117–1132. Yamanaka, T., Miyabe, S., Sawai, Y., Shimoyama, S., 2010. Geochemical and diatom evidence of transition from freshwater to marine environments in the Aira Caldera and Kagoshima next term Bay, Japan, during post-glacial sea-level rise. Journal of Asian Earth Sciences 39 (5), 386–395. Yamanaka, T., Maeto, K., Akashi, H., Ishibashi, J., Miyoshi, Y., Okamura, K., Noguchi, T., Kuwahara, Y., Toki, T., Tsunogai, U., Ura, T., Nakatani, T., Maki, T., Kubokawa, K., Chiba, H., 2013. Shallow submarine hydrothermal activity with significant contribution of magmatic water producing talc chimneys in the Wakamiko Crater of Kagoshima Bay, southern Kyushu, Japan. Journal of Volcanology and Geothermal Research 258, 74–84. Yeh, H.W., 1980. D/H ratios and late-stage dehydration of shales during burial. Geochimica et Cosmochimica Acta 44, 341–352. Yokoyama, S., 2003. Geomorphology of “Shirasu” IgnimbriteFirst ed. Kokon Shoin, Tokyo (in Japanese). Zierenberg, R.A., Shanks, W.C., 1994. Geologic, hydrothermal, and biologic studies at Escanaba Trough, Gorda Ridge, offshore northern California. In: Morton, J.L., Zierenberg, R.A., Reiss, C.A. (Eds.), Chapter 14 Sediment Alteration Associated with Massive Sulfide Formation in Escanaba Trough, Gorda Ridge: The Importance of Seawater Mixing and Magnesium Metasomatism: U. S. Geological Survey Bulletin, 2022, pp. 257–278.