Kawagucci et al. Progress in Earth and Planetary Science
https://doi.org/10.1186/s40645-018-0232-3(2018) 5:74

RESEARCH ARTICLE

Progress in Earth andPlanetary Science

Open Access

Cool, alkaline serpentinite formation fluid
regime with scarce microbial habitability
and possible abiotic synthesis beneath theSouth Chamorro SeamountShinsuke Kawagucci1,5* , Junichi Miyazaki1, Yuki Morono2, Jeff S. Seewald3, C. Geoff Wheat4 and Ken Takai1

Abstract
South Chamorro Seamount (SCS) is a blueschist-bearing serpentinite mud volcano in the Mariana forearc. Previous
scientific drilling conducted at SCS revealed highly alkaline, sulfate-rich formation fluids resulting from slab-derived
fluid upwelling combined with serpentinization both beneath and within the seamount. In the present study, atime-series of ROV dives spanning 1000 days was conducted to collect discharging alkaline fluids from the cased
Ocean Drilling Program (ODP) Hole 1200C (hereafter the CORK fluid). The CORK fluids were analyzed for chemicalcompositions (including dissolved gas) and microbial community composition/function. Compared to the ODP
porewater, the CORK fluids were generally identical in concentration of major ions, with the exception of significant
sulfate depletion and enrichment in sulfide, alkalinity, and methane. Microbiological analyses of the CORK fluids
revealed little biomass and functional activity, despite habitable temperature conditions. The post-drilling sulfate
depletion is likely attributable to sulfate reduction coupled with oxidation of methane (and hydrogen), probably
triggered by the drilling and casing operations. Multiple lines of evidence suggest that abiotic organic synthesis
associated with serpentinization is the most plausible source of the abundant methane in the CORK fluid. The SCS
formation fluid regime presented here may represent the first example on Earth where abiotic syntheses areconspicuous with little biotic processes, despite a condition with sufficient bioavailable energy potentials andtemperatures within the habitable range.
Keywords: Forearc serpentinite mud volcano, South Chamorro Seamount, Limit of biosphere, Present-days’chemical evolution, Radio-isotope-tracer carbon assimilation estimation

Introduction
Serpentinization of ultramafic rocks that are the major
components of the oceanic lithosphere produces highly reductive fluids, in which transformation of inorganic carbon
species to organic matter occur (e.g., McCollom 2013).
Serpentinization-associated geofluid systems have beenrecognized as modern analogs of plausible settings for theorigin of primordial life (e.g., Russell et al. 2014) as well as
ancient microbial communities (e.g., Takai et al. 2006; Ueda
* Correspondence: kawagucci@jamstec.go.jp
1
Department of Subsurface Geobiological Analysis and Research (D-SUGAR),
Japan Agency for Marine-Earth Science and Technology (JAMSTEC), 2-15
Natsushima-cho, Yokosuka 237-0061, Japan
5
Institute of Geochemistry and Petrology, ETH Zürich, Zürich, Switzerland
Full list of author information is available at the end of the article

et al. 2016). Serpentinization-associated geofluid systems,
typically characterized by an abundance of H2 and/or highly
alkaline fluids, have been discovered at a range of tectonic/
geological settings on the modern Earth where ultramafic
rocks encounter water circulation. The most notable examples include slow-spreading mid-oceanic ridges (Charlou etal. 2002; Konn et al. 2015), ocean core complexs (Kelley etal. 2001; Früh-Green et al. 2003), ophiolites (Barnes et al.
1978; Schrenk et al. 2013), coastal springs (Barnes et al.
1967), onshore volcanic hot springs (Homma andTsukahara, 2008; Suda et al. 2014), inner trench slope along
the southern Mariana forearc (Ohara et al. 2012; Okumura
et al. 2016a; Onishi et al. 2018), and forearc serpentinite
seamounts (Mottl et al. 2004; Fryer, 2011).

© The Author(s). 2018 Open Access This article is distributed under the terms of the Creative Commons Attribution 4.0
International License (http://creativecommons.org/licenses/by/4.0/), which permits unrestricted use, distribution, andreproduction in any medium, provided you give appropriate credit to the original author(s) and the source, provide a link to
the Creative Commons license, and indicate if changes were made.

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Forearc serpentinite mud volcanoes are formed by extensive serpentinization of the mantle, triggered by the injectionof slab-derived fluids into the overlying mantle wedge in a
subduction setting. These serpentinization products ascend
to form the largest mud volcanoes on Earth (Fryer, 2011).
Deep-sea serpentinite mud volcanoes are outstanding settings, compared to other serpentinization-associated geofluid
systems (Table 1), for studying abiotic organic syntheses (i.e.,present-days’ chemical evolution) and for finding the boundary between microbiologically habitable and uninhabitablezones (i.e., limit of life and the biosphere) in natural environments of the Earth. For example, the impermeable mud volcano allows for decreased entrainment of photosynthesisderived organic matter and seawater-derived microbial
populations into the internal part of the volcano wherepristine mud and fluid ascend from the deep. As such, if we
can obtain these pristine fluids contributing to the formationof serpentinite mud volcanoes, then they can be used to address the unique geochemical and microbiological processes
associated with these volcanoes.
The South Chamorro Seamount (SCS) is located at thesouthern end of a serpentinite mud volcano chain in the
Mariana forearc, where the Pacific plate subducts beneaththe Philippine Sea plate (Fig. 1a). To date, the SCS is the
only known location of active blueschist-bearing serpentinite mud volcanism that hosts lush deep-sea chemosynthetic
communities sustained by the serpentinite fluid seepage(Fryer and Mottl, 1997). The Ocean Drilling Program(ODP) Leg 195 drilled the summit of the SCS (Site 1200) in
2001 and obtained core samples in order to examine

Page 2 of 20

geological and biogeochemical processes associated with
subduction (Shipboard Sci. Party 2002). The serpentinite
mud porewater, with the exception of taken near the seafloor, exhibited a highly basic pH of 12.5 and had a composition rich in dissolved sulfate and carbonate. The
alkaline porewater chemistry is indicative of slab-derived
fluids that had reacted with mantle rocks when ascending
through the overlying plate (Fig. 1b), but few studies haveinvestigated the origin of the abundant sulfate in the deep
(Mottl et al. 2003, 2004; Hulme et al. 2010). In addition,
porewater samples from shallow zones ~ 20 mbsf exhibited
sulfate depletion and high alkalinity. The chemistry of theseshallow fluids has been interpreted to reflect the oxidationof CH4 (and H2) derived from the deep by a shallow, subseafloor microbial community including sulfate-reducers(Mottl et al. 2003; Aoyama et al. 2018). However, the gas
composition of the porewater could not be accurately determined, since the standard ODP sampling retrieval, handling, and on-board procedures resulted in a significant lossof dissolved gases. Microbiological analyses including
phospholipid analysis (Mottl et al. 2003), cell count, andcultivation attempts (Takai et al. 2005) were carried out,
but the community structure and potential microbial functions could not be determined.
A cased borehole observatory, the Circulation Obviation
Retrofit Kit (CORK), was deployed during the ODP Leg 195
at Hole 1200C on the SCS (Shipboard Sci. Party 2002;
Wheat et al. 2008). When the observatory seal was opened
on 20 March 2003, 2 years after its installation,
over-pressurized fluids discharged from the CORK outlet

Table 1 Deep-sea serpentinization-associated geofluid fields
Name of field

South Chamorro Seamount (SCS)

Shinkai Seep Field

Lost City

Rainbow

Location

13°47 ′N–146°00′E

11°40′N–140°03′E

30°07′N–42°07′W

36°14′N–33°54′W

Tectonics

Forearc seamount chain

Non-accretionary trench slope

Oceanic core complex, MAR

Non-transform offset, MAR

Driving force

Serpentine mud volcanism(Hydrothermalism?)Hydrothermalism

Hydrothermalism

Fluid source

Slab-derived fluid

Seawater penetration

Seawater penetration

Seawater penetration

Approach

ODP, CORK, Dive

Dive

Dive, IODP

Dive

Depth (m)

2960

5800

700–800

2300

T max

2

–

> 91

365

pH

12.5

(Brucite chimney)

10.7

2.8

H2 (mM)

0.01 (possibly < 40)

–

15

16

CH4 (mM)

37 (possibly < 65)

–

2

2.5

Note

The strongest alkaline
deep-sea geofluid
system so far identified.
See text for details.

Active alkaline fluid seepingfrom brucite chimney
surrounded by macrofaunacolony.

Hot alkaline fluid discharging
from tall carbonate chimney,
only identified on the modern
seafloor at this moment.

Black smoker acidic fluid
discharging from sulfide
chimney with abundant
H2. Similar vent sites havealready been found on
the modrern seafloor.

Ref

See text

Ohara et al. 2012

Kelley et al. 2001

Charlou et al. 2002

Stern et al. 2014

Früh-Green et al. 2003

Okumura et al., 2016a
Onishi et al. 2018

Proskurowski et al. 2008

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Page 3 of 20baFig. 1 Topography and schematic illustration of the Mariana forearc and the SCS. a Seafloor topography of the Mariana forearc region. b
Schematic illustration of the ODP Hole 1200C focused in the present study

pipe (this fluid is referred to as the ‘CORK fluid’ hereafter) ata flow rate of ~ 0.2 L/s. The endmember chemical composition of the CORK fluid, estimated from two fluid samples
diluted with entrained ambient seawater, was generally similar to those from the porewater samples from deeper zonesof the ODP core (Wheat et al. 2008). Although the deepporewater samples obtained during the ODP Leg 195 in
2001 contained abundant sulfate and negligible sulfide, sulfide was present in the highly alkaline (pH 12.4) CORK fluid.
This difference implies that partial reduction of sulfate to sulfide could occur within the SCS mud volcano before the formation fluid discharged through the cased borehole (Wheat
et al. 2008). Discharge from the open borehole continued for
37 days, until the observatory was resealed on 28 April 2003.
During the 2003 campaign, however, no gas-tight fluid sampling was conducted despite gas species such as H2, CO,
CH4, and other hydrocarbons are key chemical clues for investigating both abiotic organic synthesis and microbiologicalhabitability/function in an active serpentinite mud volcano.

The present study presents results from four subsequent sample collection campaigns utilizing remotely
operated vehicle (ROV) dives with gas-tight fluid samplers (Seewald et al. 2001; Saegusa et al. 2006) andother sampling tools conducted between 2009 and 2012
to collect fresh fluids discharging from the cased borehole of Hole 1200C. The chemistry of the CORK fluid,particularly dissolved gas species, as well as the microbial community composition and function were determined. Carbon and sulfur compositions in the CORK
fluids were altered from those from the ODP porewater
samples, probably due to changes in the subseafloorhydrological field before and after the scientific drilling
event. Analyses of the CORK fluids and the ODP porewater suggested that the sulfate-methane couple likely
co-exist in the native SCS formation fluid. However,
low biomass and activity were observed environments
where physical and chemical conditions, particularly
the temperature (~ 3 °C) and bioavailable energy

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

potentials, were suitable for supporting a prosperous
microbial community. Multiple lines of evidence, suchas low microbial biomass and activity, point to abiotic
hydrocarbon synthesis associated with serpentinization
occurring beneath the SCS.

Methods/experimental
Site description

In 2001, a series of boreholes were drilled near the seepage sites at the summit of the SCS (ODP Site 1200)(Shipboard Sci. Party 2002). The ODP Hole 1200C was
drilled to 266 mbsf prior to casing and the installation ofa CORK device. The final configuration was placed at
the base of the casing at 202.8 mbsf, providing ascreened section from 148.8 mbsf to 202.3 mbsf (Fig. 1b).
On 20 March 2003, the CORK monitoring instruments were recovered using the ROV Jason II, andon 28 April 2003 the top of the CORK was resealed
by a steel housing holding a downhole pressure sensor (Wheat et al. 2008).
Sampling and analysis
Fluid sampling

CORK fluids were collected for both chemical characterization
and analyses of microbial biomass/function during
four cruises of the R/V Natsushima with ROVHyper-Dolphin in January 2009 (NT09-01, HPD#941-947),
May–June 2009 (NT09-07, HPD#1007-1010), February
2012 (NT12-04, HPD#1349-1350), and September 2012
(NT12-23, HPD#1433-1434). On 20 January 2009
(HPD#941), the valve on the original-type CORK was
opened. The discharging fluid (#941) appeared blackish incoloration, probably due to suspension of sulfide minerals
that had accumulated during the past 6 years. The likely
source for this blackish fluid is the discharging fluid that occupied the pipe casing when the valve was closed in 2003.
After initial samples were collected, the steel plug was replaced with a polyvinyl chloride (PVC) tube and a ballvalve. This PVC tube was inserted past the latch holes in
the CORK body, sealing the borehole from the bottom seawater. Thus, fluids collected from the top of the CORK canbe considered pristine CORK fluids with minimal or no entrainment of the bottom seawater (Wheat et al. 2008). Fluid
collected during subsequent dives appeared to be colorlessand clear at the mouth of the CORK but the coloration
changed to whitish after venting, most likely due to the precipitation of brucite from seawater-derived Mg2+ and
CORK fluid-derived OH−.
CORK fluids were collected using several types of fluid
samplers, including WHATS, Bag, and Isobaric Gas-Tight(IGT) samplers (Seewald et al. 2001; Saegusa et al. 2006).
The whitish fluids were also collected into pre-evacuatedstainless-steel bottles with radio-isotope-labeled tracers
for metabolic activity measurements (see below). The inlet

Page 4 of 20

nozzles for WHATS and Bag samplers were inserted as
deep as possible into the borehole in order to minimize
seawater contamination. For fluids collected using IGT
samplers, a polycarbonate reservoir was attached to the
CORK ball valve. The reservoir was flushed with several
volumes of CORK fluid before the IGT sampler snorkel
was inserted into the reservoir through a rubber diaphragm to collect pristine CORK fluid without seawater
entrainment.
A special sampling device, named the “Deep Sampler,” wasalso used during cruises NT12-04 and NT12-23 to collect
fluid within the borehole at subseafloor depths where thescreened casing is positioned (Fig. 2). The Deep Sampler
consisted of a 100 mL gas-tight and pressure-preserved titanium bottle and a valve connected to a winch on the seafloor with a rope. The open titanium bottle-chamber-valve
units were lowered to a depth of 179 mbsf within the depth
range of the screens, before a messenger weight was dropped
from the seafloor to slide the bottle unit to thechamber-valve unit for fluid sampling (a few seconds to 10 stime) and in the end sealing the fluid in two of the chambers(Fig. 2). The temperature was monitored during fluid sampling, which increased from ~ 1.7 °C at the seafloor to 3.3 °C
at 170 m below the seafloor (thermal gradient of 9.4 °C/km).
Samples collected by the Deep Sampler and the other samplers did not show significant differences in their chemicalcomposition, implying that fluid alteration and seawater contamination during the 12 h ascent through the cased hole
(Wheat et al. 2008) was negligible.
Geochemistry

Immediately upon recovery of the ROV, fluid samples were
subsampled for liquid and gas analyses. For liquid phase analysis, the fluid sample was withdrawn into disposable plastic
syringes for filtration, and then introduced into acid-washed
HDPE bottles and glass vials. Shipboard measurements included pH, alkalinity, chlorinity, and hydrogen sulfide.
Shore-based analyses with devices such as ion chromatography (IC), inductively coupled plasma atomic emission
spectrometry (ICP-AES), and ICP mass spectrometry(ICP-MS) were conducted using techniques similar to thosecarried out for the ODP 1200 porewater samples (e.g., Mottl
et al. 2003, 2004), thus minimizing analytical biases among
these studies. Dissolved gas extraction from fluid samples
collected with the WHATS sampler was conducted using
previously reported techniques (e.g., Kawagucci et al. 2016;
Miyazaki et al. 2017). To extract the total inorganic carbon(mainly CO32−) as CO2 for abundance and isotope analyses,
the fluid sample was acidified with ~ 3 g of solid sulfamicacid (H3NSO3) in the extraction line. While the total gas
content was determined barometrically during the gas extraction, the relative abundance of H2, CH4, CO, CO2, C2H6,
and C3H8 in the whole gas phase were determined by gas
chromatography using a helium ionization detector. The

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

aPage 5 of 20bcFig. 2 A schematic illustration and photographs of the “Deep Sampler” used for CORK fluid sampling. a The open titanium bottle, chamber, andvalve units were lowered to the screened depths (170 mbsf) through the cased hole. Then, a messenger weight was dropped from the seafloor
along the rope, resulting in closing the bottle unit to the chamber-valve unit for fluid sampling. b A photograph showing the Deep Sampler
being prepared in the payload of the ROV Hyper-Dolphin. c A photograph showing the on-site operation of the Deep Sampler from 170 mbsf bymanipulation of reeling in the support line

stable isotopic compositions of H2, CH4, CO2, and C2H6
were determined using the continuous flow method with
DELTA XP Advantage and/or MAT253 Isotope Ratio Mass
Spectrometer (Thermo Fisher Scientific) in JAMSTEC in accordance with previously published techniques (e.g., Okumura et al. 2016b; Shibuya et al. 2017). In the present study,
isotope ratio and isotope fractionation between two molecules is represented by conventional δ and α notations as
δ13CA = [(13CA/12CA)/(13CSTD/12CSTD) − 1] (presented in permil scale) and αCA − B = (δ13CA + 1)/(δ13CB + 1). Samples collected using IGT samplers were analyzed at sea for dissolved
H2 and CH4 using gas chromatography with thermal conductivity detection, following headspace extraction.
Microbiology

During cruises NT09-01 and NT09-07, microbial cells in
subsampled fluids (110–370 mL) were collected
on-board the ship onto 0.2 μm pore 47 mm polycarbonate membranes (K020A047A, ADVANTEC, TokyoJapan) placed in line holder (KS-47, ADVANTEC, TokyoJapan). The cells were fixed by placing a fixation solution(2% paraformaldehyde-PBS solution) onto the membranes and kept overnight at 4 °C. Then, the membrane
was washed by PBS, partially dried by placing on filter
paper, and kept at − 20 °C for later cell counting. The
fluid sample taken on the cruise NT12-23 was immediately fixed by adding 1/5 volume of 10% Formalin Neutral Buffer Solution (Wako, Japan) and stored at 4 °C
until further processing. Since white precipitates formed
in fixed fluid samplers during storage at 4 °C, separationof cells was conducted according to Morono et al.
(2013), with some modifications. In brief, for 3–15 mL
of fixed solution, 1/10 volume each of detergent mix(100 mM EDTA, 100 mM sodium pyrophosphate, 1%[v/v] Tween 80) and methanol were added and treatedin an ultrasonic bath (Bioruptor UCD-250, Cosmo Bio,
Tokyo Japan) at 160 W with cycles of 1 min ON and
1 min OFF for 20 min. The treated samples were thencarefully layered onto step gradient heavy density solution composed of 67% (w/w) sodium polytungstate(SPT1, SOMETU, Berlin, Germany), 80%, 50%, and 30%(w/v) Nycodenz and centrifuged for 120 min at 10,000×g

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

at 25 °C. Light density upper layers were recovered, and
10% of the suspension was passed through a 0.2 μm pore
25 mm polycarbonate membrane to trap microbial cells.
The filters were stored at − 20 °C until counting. From
Bag sampler (NT09-01) and Deep Sampler (NT12-23),
about 10 L and 40 mL of fluids, respectively, were filtered with 0.22 μm-pore-sized cellulose acetate filters tocollect microbial cells. Filters were stored at − 80 °C for
microbial 16S rRNA gene clone sequencing analysis.
Microbial cell densities in CORK fluid samples were
determined using a semi-automated cell counting system(Morono et al. 2009). In brief, trapped and fixed cells
collected on a polycarbonate membrane were stainedwith SYBR green I staining solution (1/40 [v/v] SYBRGreen I in TE buffer). The stained filter was washedwith 1 ml of TE buffer and then mounted on a glass
microscope slide with 3–5 μL of mounting solution (2:1
mixture of VECTASHIELD mounting medium H-1000
and TE buffer). Microscopic fluorescence image acquisition (at 525/36 nm [center wavelength/bandwidth] and605/52 nm by 490 nm excitation) was performed automatically using a fluorescence microscope equipped with
an automatic slide handler (Morono and Inagaki, 2010).
The resulting images were analyzed using the macro of
the Metamorph software (Molecular Devices, CA, USA)to identify and count microbial cells on the membrane.
Phylotype composition of 16S rRNA genes were determined using clone sequencing in the laboratory based ona previously published methodology (Nunoura et al. 2012)
with some modifications. DNA assemblages were extracted from microbial cells on filters using the UltraClean
Soil DNA Isolation Kit (MO BIO Laboratories, Carlsbad,
CA, USA). Amplification of 16S rRNA genes from theextracted DNA assemblage was accomplished by a
Polymerase Chain Reaction (PCR) method with LA Taq
polymerase (Takara Bio, Otsu, Japan) using a primer set of
530F/907R. The amplified 16S rRNA gene fragments wereapplied to electrophoresis and cleaned up with a QIA
quick Gel extraction Kit (QIAGEN, Valencia, CA, USA).
The 16S rRNA gene clone libraries were constructedusing the TA cloning kit (Invitrogen, Carlsbad, CA, USA).
The nucleotide sequences of randomly selected clones
were determined with a BigDye Terminator version 3.1 sequencing kit (Applied Biosystems, Foster City, CA, USA)using M13M4 (5′-GTTTTCCCAGTCACGAC-3′) andM13RV (5′-TGTGGAATTGTGAGCGG-3′) primers on
an ABI PRISM 3730xl Genetic analyzer (Applied Biosystems, Foster City, CA, USA). The sequences were assembled with Sequencher Ver. 4.8 software (Gene CodeCorporation, Ann Arbor, MI, USA). The most similar reference sequences were retrieved from RDP and the
GenBank database. The nucleotide sequences obtained
were deposited to DDBJ (DNA data bank of Japan) withaccession numbers LC279286-LC279362.

Page 6 of 20

Radio isotope tracer experiment to estimate potential
metabolic rates

Potential activities of microbial assimilations of bicarbonate, methane, carbon monoxide, formate, acetate,
and leucine (autotrophic, methanotrophic, carboxydotrophic, and heterotrophic activities, respectively) were
estimated by radio isotope tracer experiment using
CORK fluid samples obtained during the expeditions
NT09-01 and NT09-07. The CORK fluid was collected
in multiple evacuated 150 mL stainless-steel bottles,
each connected to a three-way valve and a sampling
nozzle. All surfaces of the bottle, valve, and nozzle weretreated with SilicoNert coating (Silcotek Corporation,
Bellefonte, PA, USA) to avoid direct contact between thefluid and stainless steel. The sampling was conducted in
duplicates for each of the radio isotope tracers. Prior to
the seafloor sampling by the ROV, each of the14
14radio-isotope-tracers(NaH14CO3,
CH4,
CO,
14
14
14
14
H COONa, CH3 COONa, C-leucine) was placed in
two bottles (with and without formaldehyde at a final
concentration of 4%) as solution (NaH14CO3,
H14COONa, 14CH314COONa, 14C-leusin and formaldehyde) or gas (14CH4 and 14CO) in sealed glass ampoules.
The final radioactivity level was 0.25 MBq per bottle. In
order to assess the methanotrophic activity, three treatments were used: (i) only 14CH4 added, (ii) 14CH4 andO2 (a final concentration of 40 μM O2) added, and (iii)14
CH4 and Na2SO4 and Na2S (final concentrations of
10 mM and 0.5 mM, respectively) added. The additional
O2 or Na2SO4 and Na2S were introduced either as a gas
or as a solution sealed in a glass ampoule. The final concentrations of 14C-labeled tracers in the bottles were estimated to be about 0.1 mM, 0.4 mM, 0.4 mM, 0.1 mM,
0.1 mM, and 0.1 mM for bicarbonate, methane, carbon
monoxide, formate, acetate, and leucine, respectively.
After placing the radio isotope tracer into the bottle withthree stainless-steel balls, the bottle was evacuated witha vacuum pump. The bottle was then sealed and shaken,
which resulted in the stainless-steel balls rapturing theglass ampoules containing the radio-isotope-tracers andformaldehyde. At the seafloor, the sampling nozzle of
the radio-isotope-tracer sampler was placed in the discharging CORK fluid and sufficiently flushed. Fluid sample was collected by opening and closing the valve with
the ROV manipulator. Immediately after recovery of the
ROV, sample bottles were incubated at 4 °C for 49–
123 days until measurement in a shore-based laboratory.
The bottles containing 4% formaldehyde were used as
negative controls (i.e., with no microbial activity).
As a positive control for methane assimilation, stationary phase cells of a gammaproteobacterial methanotroph,
Methylomarinum vadi, (Hirayama et al. 2013) were incubated in its medium (Hirayama et al. 2007) at 37 °C for
1 day under the same substrate condition (0.4 mM) as theKawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

aerobic methanotrophic activity measurement, as described above. The cell density of M. vadi was adjusted tobe 2.0 × 106 cells in the bottle (1.3 × 104 cells/mL). As apositive control for bicarbonate assimilation, natural
deep-sea diffusing hydrothermal fluid (average fluid
temperature of 57 °C) from the Iheya North field (Takai etal. 2004), sampled in July–August 2010 using the same
method described in previous sections of the present
study, was incubated at 60 °C for 5 days under the same
substrate condition (10 mM). The cell density of microbialcommunity in the CORK fluid was 1.0 × 102 cells/mL,
thus each incubation bottle contained 1.5 × 104 cells.
CORK fluid samples (approx. 150 mL) were incubatedwith radio-isotope-tracers for 49–123 days at 4 °C andthen depressurized and filtered with a 0.22 μm-pore-sizedcellulose acetate filters. Each filter was fixed with 10 mL
of 4% (w/v) paraformaldehyde/PBS solution (pH 7.2) for
30 min, then thoroughly rinsed with filter-sterilized
0.1 M HCl in 3% (w/v) NaCl solution to remove carbonate precipitates. The filter was slurried in 20 mL ofHionic Flour cocktail (PerkinElmer, Waltham, MA,
USA), and the radioactivity measured using a Tri Carb
2910TR liquid scintillation counter (PerkinElmer, Waltham,
MA, USA). The assimilation rates of the tracers were calculated from concentrations and radio-isotope ratios given as
the substrate in the initial state and the radio-isotope abundance of the filter. Blanks were evaluated by non-labeled
batch incubation and used for calibration.

Results
pH, chlorinity, and major and minor ionsConcentrations of major and minor ions in the fluids
were nearly identical in all samples, regardless of thesampling device used for collection (Additional file 1:
Table S1). The chemistry of conservative components in
CORK fluids were quite similar to those of deep (>20 mbsf ) porewater samples (DPW) obtained from the
ODP Holes 1200 (Additional file 1: Table S1). For example, the fluid chlorinity (503–510 mmol/kg) of CORK
fluids was distinguishable from seawater level(540 mmol/kg) and identical to those of the DPW (510
± 5 mmol/kg) (Fig. 3a). CORK fluids ranged from 12.1 to12.3 in pH values, which were slightly lower than thehighest pH of 12.5 ± 0.1 in the DPW (Fig. 3b). Slight differences between CORK fluids and the DPW in alkali
metals (K, Li, Rb) (Additional file 1: Table S1) have beeninterpreted to be due to sampling artifacts during thecore recovery and porewater extraction of the ODP serpentinite samples (de Lange et al. 1992; Wheat et al.
2008). Alkali earth metals (Mg, Ca, Sr) were depleted in
the CORK fluids, as well as the DPWs (Fig. 3c). Seawater
entrainment during the seafloor sampling cannot be
evaluated precisely, because any entrainment of Mg
from seawater would be precipitated immediately as

Page 7 of 20

brucite. Nevertheless, the nearly identical and high concentrations of B, for example (Fig. 3d), are consistent
with a scenario where mixing with ambient seawater
during sample collection was minimal. Combined, thesedata are consistent with the CORK fluids being derived
from a deep-sourced formation fluid that flowed into theborehole through screens located at a depth range of
149 and 202 mbsf.
Sulfate, alkalinity, sulfide, and ammonium

Sulfate concentrations in the discharging CORK fluids
were 0.4–1.7 mM, significantly lower than the concentration of 28 mmol/kg in the DPW and the seawater(Fig. 3e). The alkalinities of these fluids were muchhigher (119–126 meq/kg) than in the DPW (62 meq/kg)(Fig. 3f ) and comparable to the value estimated from theprevious CORK fluid observation (104 meq/kg) (Wheat
et al. 2008). Both sulfate concentration and alkalinity in
the CORK fluid from Dive #941 were in-between valuesof the DPW and other CORK fluids. These results point
to a stoichiometry of 2 meq increase of alkalinity with
1 mmol decrease of sulfate in the CORK fluids whencompared to the DPW (Fig. 4a). Dissolved reduced sulfur (sulfide) in CORK fluids was significant despite negligible sulfide concentrations detected in the DPW andshowed a temporal decreasing trend: ~ 24 mM in
10 days, 21 mM in 4 months, and ~ 10 mM in morethan 36 months (Fig. 3g). The sum of sulfate and sulfide
concentrations in the CORK fluids were not higher than
28 mM, suggesting quantitative conversion in addition
to the removal of dissolved sulfur to the solid phase(Fig. 4b). Ammonium concentrations were identical at
0.2–0.3 mM among all CORK fluids and the DPW(Additional file 1: Table S1). The highly variable alkalinity with the constantly low ammonium concentrations
implies that the alkalinity variation do not result fromthe decomposition of sedimentary Redfield organic matter (Redfield, 1958) but from that of N-depleted organiccarbon such as CH4.
Gas composition and stable isotopes

Methane concentrations in the CORK fluids were variable
between 5.5–6.1 mM in 10 days, 9.2–37.8 mM in 4 months,
and 16.1–37.4 mM in more than 36 months after the valve
was opened, respectively (Fig. 3h). CH4 concentrations in
CORK fluids were significantly higher than those observedin the DPW (< 2 mM). The differences strongly suggest that
CH4 was degassed from the DPW during core recovery andsample processing during the ODP Leg 195 (Mottl et al.
2003). This claim is consistent with a solubility-controlled
CH4 concentration of ~ 2 mM at room pressure andtemperature (Wiesenburg and Guinasso jr, 1979). The CH4
concentrations in CORK fluids are the highest among any
deep-sea serpentinization-associated geofluid systems

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Page 8 of 20

a

b

c

d

e

f

g

h

Fig. 3 Time-series of fluid chemistry since the valve opening. Light and dark gray bars represent the deep part of the ODP porewater at Site 1200
(DPW) (Mottl et al. 2003) and ambient seawater (ASW) composition, respectively. Each symbol represents a type of fluid sampler used: WHATS(circle), Bag (hexagon), IGT (diamond), and Deep Sampler (square). Time (x axis) is in logarithmic scale

reported to date (< 2 mM) (Charlou et al. 2002; Kumagai etal. 2008; Schrenk et al. 2013; Konn et al. 2015; McDermott
et al. 2015; Seyfried et al. 2011). Ethane concentrations werealso variable between 0.01 and 0.23 mM. Relative
abundances of CH4 to C2H6 (C1/C2) were ~ 27 in 1 day,

216–3178 in 10 days to 4 months, and 231–285 in ~36 months after the valve opening. Molecular hydrogenconcentrations in the CORK fluids ranged between 0.003
and 0.107 mM and were the lowest among values repoted
from deep-sea serpentinization-associated geofluid systems

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Page 9 of 20

a

b

c

d

Fig. 4 Relationships between sulfate and alkalinity, sulfide, CH4, and pH. Filled triangles represent the CORK fluids from the present study,
classified by the number of days after the valve was opened. Filled and open circles represent the CORK fluid observed in 2003 and DPW,
respectively, shown for comparison. Light gray zones represent the estimated endmember formation fluid composition of the ODP porewater(DPW) or a solubility of CH4 hydrate clathrate under the SCS condition (only shown in panel c). Broken lines in panels (a), (c), and (d) represent
ideal changes in the chemical composition based on selected reactions. The dark gray zone in panel (c) represents possible CH4 concentrationsof the SCS formation fluid (FF) without post-drilling alteration. The solid curve in panel (b) represents sulfate consumption through CH4-SO4
reaction from the porewater-based endmember. Broken curves in panel (b) represent sulfate consumption with 10, 40, and > 112 mM of H2, in
addition to the CH4-SO4 reaction

to date (> 1 mM) (Charlou et al. 2002; Kumagai et al. 2008;
Schrenk et al. 2013; Konn et al. 2015; McDermott et al.
2015; Seyfried et al. 2011). Relative abundances of CH4 to H2(CH4/H2) ranged from 333 to 1350 for the CORK fluids.
This CH4/H2 ratio is higher than those of seafloor geofluids
associated with serpentinization in other geologic settings(< 1). Total dissolved inorganic carbon (DIC), extracted
and quantified as CO2, ranged from 17 to 51 mM, whichis consistent with the estimated carbonate alkalinity of theDPW (45 meq/kg: Mottl et al. 2003). The potentially insufficient acidification of DIC samples prior to the gas
analysis may result in the underestimation of DIC concentrations in the CORK fluids.
Carbon and hydrogen isotopic ratios of methane(δ13CC1 and δDC1) and ethane (δ13CC2 and δDC2) were
uniform among the CORK fluids, with values of − 37‰, −117‰, − 11‰, and − 86‰, respectively, regardless of concentration (Additional file 1: Table S1). The δ13CC1 valueof CORK fluids is significantly lower than otherserpentinization-associated hydrothermal fluids reported
to date (> − 18‰) (Fig. 5a) (Charlou et al. 2002; Kumagai

et al. 2008; Schrenk et al. 2013; Konn et al. 2015; McDermott et al. 2015). The higher δ13CC1 value in the shallow
ODP porewater observed previously (− 11 ± 5‰ at
15 mbsf: Mottl et al. 2003) is significantly different fromthe CORK fluid and is likely attributed to 13C enrichment
in the remnant CH4 due to microbial methanotrophic activity in the fluid-seawater interface. The δ13CC2 value of
the CORK fluids falls into the range of δ13CC2 values observed in serpentinization-associated hydrothermal fluids(− 14‰ to − 2‰) (Konn et al. 2015). The δDC2 value of
the CORK fluids (− 86‰) is higher than that of the Lost
City field (~ − 150‰) (Fig. 5b). Total DIC of the CORK
fluids showed δ13CDIC values of − 2.9‰ to + 1.4‰. The
δ13CDIC values close to + 0‰ suggest that the DIC originates from sedimentary carbonate (~ + 0‰) rather thanphotosynthesis-based organic matter (~ − 25‰) (Sano andMarty 1995). Hydrogen isotope ratios of H2 (δDH2) were
− 787‰ to − 761‰, which can be attributed to H2-H2Ohydrogen isotope equilibrium at an in situ temperature of
2 °C (Horibe and Craig, 1995) when using the δDH2O
value of + 12‰ observed in the DPW (Mottl et al. 2003).

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Page 10 of 20

Fig. 5 Stable isotope composition of CH4, C2H6, and the relevantmolecules. Open and filled symbols, respectively, represent datafrom observations of hydrocarbon-enriched geofluid systems andexperiments simulating methanogenesis. Data sources are following:high-temperature hydrothermal fluids (Konn et al. 2015; Kawagucci
et al. 2016 and references therein), Lost City (Proskurowski et al.
2006; 2008), Happo (Suda et al. 2014), The Ceders (Wang et al. 2015),
Kidd Creek (Sherwood-Lollar et al. 2002), petroleum field (Burrussand Laughrey 2010), magnetite experiments (Fu et al. 2007), Fe
experiment (McCollom et al. 2010), thermogenic experiment(Seewald et al. 1994) (assuming the CH4-H2O isotope equilibrium),and microbial incubation (Okumura et al. 2016b)

Microbial cell density

Microbial cell densities of CORK fluids were estimatedby epifluorescence microscopy (Table 2). In the blackish
fluid discharging at the time when the valve was first
opened (Dive #941), we measured a microbial cell density of 2.0 × 103 cells/mL. This is much lower than thecell density in the ambient seawater (2.2 × 104 cells/mL)
sampled during the same cruise (Table 2). The microbial
cell densities in the clear CORK fluid sampled 4 days
and 4 months after the valve opening decreased to 4.4–
6.6 × 102 and 1.0 × 102 cells/mL, respectively (Table 2).
These results suggest that, before the valve opening, certain microbial communities maintained their populations and functions in some of the microhabitats in thestagnant CORK fluid of the sealed borehole and wellhead structures. Such local microbial populations in thestagnant fluid system were flushed out by the newly created discharging fluid flow after the valve was opened.
After more than 36 months, however, the microbial celldensities in the CORK fluid were 4.1–7.6 × 103 cells/mL(Table 2).
Microbial phylotype compositions

The 16S rRNA gene phylotype compositions of microbial communities in CORK fluid samples at the time thatthe valve was opened, 3 days later, and 45 months later
were examined. In total 89, 88, and 44 clones were sequenced for the microbial communities in the CORK
fluid samples at the time of valve opening, 3 days later,and 45 months later, respectively (Fig. 6). The phylotype
compositions were significantly different among the fluid
samples.
In the blackish discharging fluid (2.0 × 103 cells/mL of microbial cell density), the gamma-proteobacterial phylotypes
dominated the clone library and consisted almost entirely ofa single phylotype closely related with Methylobactor, amethanotrophic endosymbiont of Bathymodiolus mussels
living in the SCS seepage sites (Fig. 6). The alpha-proteobacterial and actinobacterial phylotypes were the second mostabundant groups in the library. Other components were
phylotypes of Betaproteobacteria and Firmicutes. In the

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Page 11 of 20

Table 2 Microbial cell density in the CORK fluid and ambient seawater samples
Cruise ID

HPD dive#

Time after valve opening

Sample type

Cell count (cells ml−1)

NT09-01

941

0 day

Bag

2000

NT09-01

943

2 days

Bag

440

943

(ambient seawater)Bag

22,000945

4 days

Bag

660

1007

5 months

Bag

101

1009

5 months

Bag

139

1349

45 months

WHATS

7600

1350

45 months

Deep sampler

4100

NT09-07

NT12-04

CORK fluid obtained 3 days after opening (4.4–6.0 × 102
cells/mL of microbial cell density), the gamma-proteobacterial phylotypes, especially the Methylobactor phylotype, significantly decreased in the relative abundance andalpha-proteobacterial phylotypes also decreased in the relative clonal abundances (Fig. 6). In contrast, phylotypes ofDeltaproteobacteria, Firmicutes (Alkaliphilius), Actinobacteria, Deinococcus-Thermus, and Archaea increased in relative abundances. Most archaeal phylotypes were related to
extremophilic members such as hyperthermophiles and extreme halophiles that do not inhabit the cold and highly alkaline serpentine formation fluid regime of the SCS (Fig. 6).
Although this simplified interpretation may lead to misunderstanding since the microbial cell density was decreased by

fluid replacement during the 3-day discharge of the CORK
fluid, phylotypes that increased in clonal abundances 3 days
after the valve opening more likely represent truly indigenous microbial populations entrained by the serpentinite formation fluid derived from the deep. In the CORK fluid
obtained by the Deep Sampler 45 months after the valve
opening (4.1–7.6 × 103 cells/mL of microbial cell density),
the microbial phylotype composition changed drastically(Fig. 6). Most of the predominant phylotypes found in thefluids were different from those sampled 3 days after thevalve was opened. At 45 months after, phylotypes of Deltaproteobacteria, Epsilonbacteria, Planktomycetes, Thermus-Deinococcus, and unclassified bacterial group dominated thelibrary. It is still unclear why the CORK fluid obtained by theFig. 6 Phylotype compositions of 16S rRNA genes in microbial communities of the CORK fluids. In total, 88, 89, and 44 clones were sequenced
for the microbial communities in the CORK fluid samples at the time of valve opening (0 day), 3 days, and 45 months after the valve was
opened, respectively

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Deep Sampler 45 months after had a higher cell density anda different microbial phylotype composition compared to theinitial phase of CORK fluids. Nevertheless, it should be notedthat the sampling scheme by the Deep Sampler was different
from those by other fluid samplers, with the Deep Sampler(5 cm in diameter with the bottle being 1 m in length) was
delivered to 170 mbsf through the cased borehole. During
the descent, the bottle could have hit the casing wall and detached accreted materials which likely contain microbial
populations. Thus, the microbial community abundance andcomposition in the fluid obtained by the Deep Sampler may
include, in addition to the microbial communities entrained
from the formation fluid (“pelagic”) flow, microbial populations specifically colonizing the steel casing pipe wall (Nakagawa et al. 2006; Nigro et al. 2012).
Microbial carbon assimilation activity

Microbial methane-, carbon monoxide-, formate-, acetate-, and leucine-assimilations (methanotrophic, carboxydotrophic, and heterotrophic activities) were examinedwith the CORK fluid samples obtained 1 to 6 days after
the valve opening, and the bicarbonate-, methane-, formate-, and acetate-assimilation activities were examinedwith the CORK fluid samples obtained 5 months after
the valve opening (Table 3). Methylomarinum vadi is amarine aerobic methanotroph optimally growing at 37 °C
with a generation time of 2 h (Hirayama et al. 2013), andthe methane assimilation rate of this methanotroph underthe maintenance state for 1-day incubation was used as
the positive control: 20 nmol/day/L with a cell density of
1.3 × 104 cells/mL (Table 3). In a similar manner, the autotrophic bicarbonate assimilation of natural deep-sea diffusing hydrothermal fluid was determined to be 2.3 nmol/
day/L with a cell density of 1.0 × 102 cells/mL (Table 3).

Page 12 of 20

These methane assimilation rates of M. vadi and the bicarbonate assimilation of the natural autotrophic microbial community in the discharging fluid are goodcomparative indexes to estimate the functions of themethanotrophic and autotrophic populations and even
the whole microbial community in the CORK fluid.
The methane assimilation rates were determined to be
0.036, 0.049 and 0.039 nmol/day/L in the CORK fluid underin situ, microaerobic and fully anaerobic conditions, respectively (Table 3). These values were significantly higher thanthose of the negative controls with the same tracers and are
attributed to activity from the microbial populations in the
CORK fluid. In addition, these values correspond to the methane assimilation activity by 2.3–3.2 × 101 cells/mL throughaerobic methanotrophy in the maintenance state. Carbon
monoxide assimilation was not detected, while heterotrophic
carbon assimilation was found to be the highest with formate(0.06 nmol/day/L) and the lowest with acetate (0.004 nmol/
day/L). These values are similar to or lower than the methane assimilation activity. In the CORK fluid obtained
5 months after the valve opening, autotrophic bicarbonate
assimilation activity was the highest (1.4 nmol/day/L), andmethanotrophic and heterotrophic (with formate and acetate) assimilations were lower at 0.2 and 0.05 nmol/day/L, respectively (Table 3). The methanotrophic assimilation was
higher than any of the methane assimilation activities using
the CORK fluid obtained 1 to 6 days after valve opening andthe heterotrophic activity was similar to values reported
previously from the CORK fluid, although the cell density in
the CORK fluid decreased from 4.4–6.6 × 102 to 1.0 × 102
cells/mL during the 5 months period. Compared to thevalues of positive controls, these autotrophic andmethanotrophic assimilation activities correspond to
7.0 × 101 cells/mL of a natural autotrophic microbial

Table 3 Radio-isotope-tracer carbon assimilation of microbial community in the CORK fluids and positive controls
Cruise ID

HPD dive#

Tracer and incubation condition

Incubation Net assimilation
time (days) (nmol day−1 L−1)

NT09-01

943

14

54

943

14

54

0.049

947

14

52

0.014

947

14

52

0.06

947

14

52

0.004

948

14

948

14

49

n.d.a

1007

14

C-CH4 (0.25 MBq per bottle and 0.4 mM)

123

0.2

1007

14

C-HCO3− (0.25 MBq per bottle and 0.1 mM) with 0.4 mM H2

123

1.4

1009

14

NT09-01

NT09-01

NT09-07

C-CH4 (0.25 MBq per bottle and 0.4 mM)
C-CH4 (0.25 MBq per bottle and 0.4 mM) with 40 μM O2
C-Leucine (0.25 MBq per bottle and 0.1 mM)
C-Formate (0.25 MBq per bottle and 0.1 mM)
C-Acetate (0.25 MBq per bottle and 0.1 mM)

C-CH4 (0.25 MBq per bottle and 0.4 mM) with 10 mM NaSO4 and 0.5 mM Na2S 49
C-CO (0.25 MBq per bottle and 0.4 mM)

C-Formate and

122

0.05

1

20

C-HCO3− (0.25 MBq per bottle and 0.1 mM)

5

2.3

14

An autotrophic community indiffusing hydrothermal fluid

14n.d. not detected

C-acetate (each of 0.25 MBq per bottle and 0.1 mM)

0.039

C-CH4 (0.25 MBq per bottle and 0.4 mM)

An aerobic methanotroph culture

a

14

0.036

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

community and 1.3 × 102 cells/mL of relatively inactive
aerobic methanotrophs. These results indicate that thesmall number of microbial cells present in the CORK fluid
obtained 5 months post-valve opening may be functionallyactive and capable of methanotrophic and autotrophic
metabolisms. Nevertheless, radio-isotope-tracer experiments clearly indicate that microbial metabolisms and activities in the CORK fluids are detectable.

Discussion
Post-drilling fluid alteration
Sulfate reduction coupled with CH4 oxidation

Except for sulfate, the consistency in the major ion fluid
chemistry between the CORK fluid and the DPW suggests that CORK fluids originate from the deep SCS formation fluid. Sulfate depletions in the CORK fluids, as
opposed to the sulfate enrichment seen in the DPW,
likely resulted from perturbation of the subseafloor
geochemical environment during drilling activity. To
estimate the abundance of reactive components (particularly CH4 and H2) in the pristine SCS formation
fluid, it is necessary to consider processes associatedwith the drilling operation that could alter the fluid
chemistry at ~ 4 °C.
A previous study (Wheat et al. 2008) pointed out thatthe methane oxidation can cause SO4 depletion in the
CORK fluid, as follows:
CH4 + OH− + SO42− → CO32− + HS− + 2H2O (R1).
R1 produces 1 mol of dissolved sulfide and two equivalents of alkalinity for each mole of OH−, CH4, and SO42−
consumed. This stoichiometry is consistent with thestoichiometry exhibited by the differences between the
DPW and CORK fluids (Figs. 3 and 4). Assuming that
R1 fully accounts for the ~ 27 mM-equivalent sulfate
consumption in the CORK fluids, the pristine SCS formation fluids would have contained CH4 at concentrations of 35–65 mM (Fig. 4c). The measured pH values of
the CORK fluids (12.1–12.3) were higher than would be
expected given the consumption of ~ 27 mM OH− and a
DPW with a pH = 12.5 ± 0.1 that is representative of the
SCS formation fluid (Fig. 4d). In addition, some of themeasured total DIC contents of the CORK fluids (17–51 mM) were also inconsistent with a DIC concentrationof ~ 50 mM expected by a combination of the ~ 27 mM
DIC production and the 23 mM DIC being originallypresent in the pristine SCS formation fluid (Mottl et al.
2003). Differences in the OH− and DIC values between
the measured and predicted fluid compositions may be
due to brucite-magnesite buffer through the following
reaction:
Mg(OH)2 + CO32− ⇔ MgCO3 + 2OH−(R2).
The maximum CH4 concentration estimated (65 mM)
seems too high but is not unrealistic because it is comparable to the solubility of CH4 clathrate hydrate under

Page 13 of 20

relevant concentration-pressure-temperature-salinity condition within the SCS (~ 60 mM at 30 MPa, 2 °C and 3.5%)(Tishchenko et al. 2005). No geophysical data, however,support the presence of layered hydrate boundary, such as
bottom simulating reflector (BSR).
Possible H2 consumption coupled with sulfate reduction

The low concentrations of H2 in the CORK fluid (< 0.1 mM)
suggest that it may be consumed by sulfate reduction in
addition to CH4. H2 oxidation coupled with sulfate reduction
progresses as follows:
4H2 þ SO4 2− →HS− þ OH− þ 3H2 O ðR3Þ:
Because R3 produces OH−, H2 consumption during sulfate reduction may contribute to measured pH values thatare higher than predicted by R1 (Fig. 4d). Assuming the
pH balance is made by R1 and R3, the native SCS formation fluid may contain as much as ~ 40 mM of H2. ThisH2 level may be realistic because 27 mM of H2 has been
observed in a seafloor serpentinization-associated hydrothermal system at > 296 °C (Konn et al. 2015) and >100 mM of H2 can be attained through serpentinization at
200–300 °C from thermodynamic calculation (McCollom
and Bach 2009). If the cool SCS formation fluid contains
such abundant H2, the reduction of CO2 to form CH4 isthermodynamically favorable, according to the following
reaction:
4H2 þ CO3 2− →CH4 þ H2 O þ 2OH− ðR4Þ:
However, abiotic CH4 formation represented by R4 isunlikely due to kinetic barriers at the low temperatures
within the SCS body (4 °C). In addition to CH4 and H2,native iron and other metals in the casing pipe may
function as reductant and cation for the dissolved sulfate
removal.
Drilling induced sulfate reduction

The discussions above suggest that abundant CH4 andSO4 coexist in the native SCS formation fluid and isconsistent with their presence in the DPW (Mottl et al.
2003). This coexistence is unusual in the subseafloor environment because sulfate reduction coupled with methane oxidation is thermodynamically favorable andusually promoted by ubiquitous microbial consortia of
methanotrophic archaea and sulfate reducing bacteria,
known as AOM (Orphan et al. 2001). In other words,
both abundant CH4 and SO4 in the native SCS formation fluid point to unusually slow kinetics of the sulfate
reduction (R1) within the native SCS body. In fact, SO4
reduction coupled with CH4 oxidation (R1) likely occurs
in the post-drilling CORK fluids, derived from

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

subseafloor regions affected by the Hole 1200C drilling,casing, and/or CORK installation.
Drilling-induced introduction of external microbialcommunities into the subsurface environment may account for the post-drilling progression of R1 due to
AOM. While the relatively impermeable serpentinite
matrix prevented access of the over-pressured aquifer tothe microbe-bearing seawater, the drilling operation andthe hole made the deep aquifer exposed to deep-sea
water and shallow subseafloor sediment containing exogenous microbial communities, including AOM consortia (Mottl et al. 2003). However, the post-drilling
microbial AOM function is likely insignificant, as suggested by multiple lines of microbiological evidence suchas small biomass, a lack of evidence for 16S rRNA gene
phylotype signatures of previously known AOM consortia, and the negligible CH4 assimilation during the 14C
experiment (see the “Results” section). Assuming thatthe dissimilatory AOM activity of consortia does not exceed 102 times higher than the anoxic methane assimilation rate, as the highest methane assimilation activity
estimated in this study is 0.0002 μmol/day/L, it can becalculated that a 27 mmol/L SO4 consumption requires> 106 days. This amount of time is much longer than thetime elapsed since drilling ceased (< 104 days). This calculation, together with other microbiological results,suggests that the microbial AOM function is not significantly involved in the post-drilling CH4-SO4 reactions,based on experimental conditions at 2 °C. Although faster microbial reaction rates may occur deeper within the
seamount where the environment is warmer, the extent
of their involvement would be still negligible. It is notedthat microbial activity at the pristine interior of the SCS
body is more stricly limited than the rate calculated
above due to their impermeability and over-pressured
aquifer, and probably less biomass.
There are no reliable data for the kinetics of R1 without
the involvement of microbial AOM. However, the sudden
and rapid progression of R1 after drilling would require
some catalytic effects associated with the drilling, casing,and CORK installation operations, other than the introduction of AOM populations. A possible catalyst introduced by drilling and casing the borehole is native ironand other metals used in the casing pipe and the CORK
structure. Transition metals such as iron and nickel generally exhibit catalytic function for redox reactions due totheir multiple valences. In addition to catalytic reactions
occurring locally on the metal surface, high electron conductivity of the elongate casing pipe allows remotely located reducing and oxidizing molecules to be reacted
through electron transfer, similar to how a battery functions. Nevertheless, the extent to which drilling and casing
operations promoted abiotic sulfate reduction coupledwith CH4 oxidation is still not entirely clear.

Page 14 of 20

Origins of CH4 and SO4 in the native serpentine formation
fluid of the SCS
Abiotic hydrocarbon synthesis associated withserpentinization—environmental factors

Although it is difficult in natural environments of themodern Earth to distinguish between abiotic synthesis of
CH4 from other CH4 sources such as microbial methanogenesis (biogenic CH4) and thermal degradation of sedimentary organic matter (thermogenic CH4) (see reviews,
Etiope and Sherwood Lollar, 2013 for example), multiplelines of geochemical and microbiological evidence suggestthat contributions of biogenic and thermogenic CH4 in
the SCS system are not signficant. During the early stageof subduction, where the temperatures are below the
known 122 °C upper limit for the growth of microbes(Takai et al. 2008), microbial methanogenesis is probably
absent or insignificant due to the abundantly dissolved
sulfate in the sediment pore water in the Pacific plate eastward from the Mariana Trench (ODP sites 800–802)(France-Lanord et al. 1992). During the later stages of subduction, temperatures exceed 122 °C, which is too highfor microbial life to survive. The same is also true for regions of extensive serpentinization within the overlying
mantle wedge. In the SCS formation fluid regime, where
temperatures again become cool and habitable, there
is little potential to produce abundant CH4 due tolow biomass and microbial function, revealed by the
45 months long observation of the CORK fluids. In
addition, the highest known pH limit for the growthof methanogens is < 10 (Methanosalsum zhilinae:
Mathrani et al. 1988), significantly lower than the pHof > 12 in the SCS fluids. The extremely low organic
contents in the SCS serpentine mud (Mottl et al.
2003) indicate that there is little potential for thermogenic hydrocarbon generation beneath the SCS.
Possibility of abiotic synthesis from thermodynamic
and kinetic viewpoints Abiotic organic synthesis, including the generation of CH4 from CO2, is generally
more favorable at lower temperatures from a thermodynamic viewpoint despite unfavorable kinetic constraints (e.g., McCollom, 2008). In the case of the SCS,
thermodynamically favorable conditions for the abiotic
synthesis of CH4 is expected through the entire formation fluid regime due to sufficiently reducing conditions
resulting from serpentinization occur at 200–300 °C(McCollom and Bach, 2009) and subsequent cooled as
fluids ascend. The ascending fluid is quenched sufficiently for microbial activity (122 °C: Takai et al. 2008)at 13 km below the seafloor, estimated from the thermal
gradient observed (9.4 °C/km). Since formation fluids
probably ascend on a geological time span, continuous
abiotic synthesis and accumulation of organic compounds is plausible, even at extremely low reaction rates.

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

The rate of abiotic hydrocarbon synthesis depends onthe carbon substrate (McCollom et al. 2010). McCollom
et al. (2010) demonstrated that CO-based synthesis isfeasible within an experimental timeframe. Even when
the CO-based abiotic synthesis is feasible under conditions of the SCS formation fluid, CO production fromCO2 represents a bottle-neck in the process due to theseveral orders of magnitude less CO compared to CO2
in the CORK fluid. Abiotic methanogenesis based onsolid-phase carbon, such as graphite derived from metamorphism and carbonate mineral deposits (Etiope andSherwood Lollar, 2013), cannot be precluded, although
the possibility cannot be evaluated here due to limited
information about the rates of such processes.
Suggestions from stable isotope signatures Stable isotope signatures should be interpreted in the context of
isotope ratios of individual molecules (Fig. 5a, b) andisotope fractionation between relevant molecules (Fig. 5c)(Etiope and Sherwood Lollar, 2013). The CH4-H2Ohydrogen isotopic fractionation factor of the CORK fluid(αHCH4-H2O) is 0.87, which is inconsistent with typicalbiogenic values of ~ 0.70 (Okumura et al. 2016b)(Fig. 5c). Thermogenic CH4 reservoirs generally have
αHCH4-H2O values corresponding to the CH4-H2O isotope equilibrium values at the formation/reservoir
temperature (e.g., Stolper et al. 2014; Wang et al. 2015).
The αHCH4-H2O and αHC2H6-H2O values of the CORK
fluids (> 0.86) correspond to those of the isotope equilibrium values at > 200 °C (Fig. 5b, c) (Horibe and Craig
1995; Reeves et al. 2012), consistent with the uppermost
temperature expected at the plate boundary and mantle
wedge beneath the SCS (~ 350 °C) (Mottl et al. 2004;
Hulme et al. 2010). This suggests that the hydrogen isotope signature of hydrocarbons recorded at > 200 °C has
not been altered after these fluids have been cooled totheir discharge temperatures. The lack of hydrogen isotope re-equilibrium between CH4 and H2O seems consistent with the half-life for equilibrium exchange, estimatedto be approximately one million years at 200 °C (Sessions
et al. 2004; Wang et al. 2015). In addition, the hydrogenisotope signature shows no evidence for a contribution ofbiogenic CH4 produced in the cool formation fluid regimeof the SCS. The H2-H2O thermometer value, indicating ~2 °C, is inconsistent with the CH4-H2O thermometer but
reasonable because of the more feasible H2-H2O equilibrium, even at lower temperatures (Valentine et al. 2004;
Proskurowski et al. 2008; Kawagucci et al. 2010; 2016; Pester et al. 2018).
The δ13C-CH4 values of about − 37‰ and the isotope
fractionation between CH4-DIC (αCCH4-DIC) of ~ 0.963
seem to be an equivocal signature for thermogenic, biogenic, or abiotic origins (Fig. 5c) (Kawagucci et al. 2013
and references therein). Hydrothermal experiments

Page 15 of 20

simulating hydrothermal hydrocarbons generation fromthe seafloor sediment at 325–400 °C show an αCCH4-CO2
as high as 0.985 (Seewald et al. 1994) (Fig. 5c). Carbon
isotope fractionation produced by microbial hydrogenotrophic methanogenesis in the seafloor sediment typically displays αCCH4-DIC of ~ 0.930 (Okumura et al. 2016band references therein). It is also known, however, thatmicrobial metabolisms sometimes cause unusual isotope
effects under extreme growth and survival conditions(e.g., Valentine et al. 2004; Penning et al. 2005; Takai etal. 2008). Microbial methanogenesis in the SCS formation fluid regime may cause an unusual carbon isotope
effect as CO32− is the dominant DIC speciation in highly
alkaline solutions.
Carbon and hydrogen isotope systematics of CORK
fluids suggest that abiotic methanogenesis is occurringat > ~ 200 °C. Hydrothermal experiments, demonstrating
CO- and CO2-based abiotic hydrocarbon syntheses, provide consistent αCCH4-CO2 values of 0.97–0.98 (Fu et al.
2007; McCollom et al. 2010), slightly higher than thatobserved in the CORK fluids (0.963). Experimental determination of CH4-CO2 isotope equilibrium at 200 °C
exhibits a αCCH4-CO2 of 0.966 (Horita, 2001), very close
to the value observed from the CORK fluid (Fig. 5c). Although carbon isotope exchange between CH4-CO2 below~ 400 °C is generally thought to become kinetically sluggish
without a catalyst, particularly in the CH4-generating direction (e.g., McCollom, 2008), the CH4-CO2 isotope equilibrium could be attained at ~ 200 °C beneath the SCS
because of thermodynamically favorable condition for CH4
generation and geological time span. For the hydrogen isotope equilibrium between hydrocarbons and H2O, terminal
locality of hydrogen atoms would be more feasible for theexchange with the surrounding hydrogen-bearing molecules, namely, H2O (Reeves et al. 2012; Wang et al.
2015). The H exchange would erase the extremely
D-depleted signature (αHCH4-H2O = ~ 0.45: McCollom
et al. 2010) exhibited by the (experimental) abiotic
CH4 generation. Carbon and hydrogen isotopic thermometers both indicating > ~ 200 °C left in the 2 °C
CORK fluid (Fig. 5c), imply that the reactions for isotope exchange are limited at lower temperatures.
Thus, the rate of abiotic synthesis would decease after
cooling the SCS formation fluid below 200 °C, yet it
must continue within the serpentinite matrix at lower
temperatures.
Slab-derived SO4 conservation through serpentinization

The abundant dissolved sulfate in the SCS formation fluid
was thought to originate from sulfate in the subducting
slab (Mottl et al. 2003). This conclusion seems inconsistent with abiotic organic synthesis via serpentinization, in
which sulfate should be reduced from a thermodynamic
viewpoint. On the other hand, closed-system experiments

Kawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

simulating serpentinization at 200–300 °C (Janecky andSeyfried Jr., 1986; Seyfried 2007) confirmed barriers to sulfate reduction during serpentinization due to rapid sulfate
precipitation as anhydrite upon initial heating of theexperiment and its subsequent metastable persistence,even in a highly reducing fluid. Accordingly, anhydrite
precipitation in the downgoing slab may facilitate sulfate
conservation through high-temperature regions of subduction and diapir. In addition, a recent quadruple sulfur
isotope study of SCS DPW and CORK fluid suggests thatthe mantle wedge represents one of the possible sourcesof sulfate in the SCS formation fluid (Aoyama et al. 2018).
Hypothetical sequence of geochemical processes in the
SCS system

We propose a hypothesis for the possible sequence of(bio) geochemical processes in the SCS serpentinite formation fluid regime, based on CORK fluid characteristics and observations from previous studies (Fig. 7). The
oligotrophic western North Pacific water provides little
sedimentary organic matter, resulting in pelagic sediments of the Pacific plate eastward from the Mariana

Page 16 of 20

Trench (ODP sites 800–802) that contain sulfate-richporewater (France-Lanord et al. 1992; Alt and Shanks
III, 2006) and low sedimentary biomass (Kellmeyer et al.,2012). Global seafloor drilling that penetrates into theupper basaltic crust (e.g., Gieskes et al. 1990, 1993;
Hensen and Wallmann, 2005) have indeed revealed dissolved sulfate that is ubiquitously abundant in the uppermost part of subducting plate, despite their great depth
and distance from the trench seafloor. At an early stage ofsubduction, some or all of the porewater sulfate is precipitated as anhydrite with requisite Ca supplied by dissolutionof sedimentary calcite or reactions with involving basaltic
crust. At the deep plate boundary region at ~ 0.6 GPa and ≤~ 350 °C (Hulme et al. 2010), metamorphic processes suchas compactive dewatering, expulsion of interlayer water, dehydration, and decarbonation forms the DIC-rich initial upwelling fluid (Mottl et al. 2004), which harbors no microbial
population due to its high temperature. The initial fluid accompanied with metamorphic minerals such as blueschistand anhydrite penetrates into the overlying mantle wedge
and induces serpentinization. Extensive serpentinization produces highly alkaline and strongly reducing H2-rich fluid

Fig. 7 Illustration for possible C-S-H reactions beneath the SCS. “Cells” represents microbial cells and “Org” represents organic matter inthe seawaterKawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Page 17 of 20(McCollom and Bach 2009), leading to the thermodynamic
conditions favorable for the abiotic organic synthesis fromslab-derived inorganic carbon. Sulfate minerals are not
reduced, due to kinetic constraints. The abiotic synthesis oforganic molecules occurs within the ascending serpentinite
mud until its temperature decreases to ~ 200 °C. Further ascend and cooling decrease the rate of abiotic synthesis whileallowing for the dissolution of anhydrite (anhydrite-derived
calcium is completely removed by abundant DIC). This sequence results in the co-existence of abiotic CH4 (andpossibly H2, C2+ hydrocarbons, and some organic
matter) and slab-derived dissolved SO4 in the SCS formation fluid. Although the abundant CH4 and SO4 exhibit sufficient bioavailable energy potentials (Nakamura and Takai2014) within the habitable temperature range (< ~ 122 °C:
Takai et al. 2008), limited microbial biomass and activity
are present because the relatively impermeable SCS prevents penetration of microbial community from ambient
seawater. As the major conduit for the mud/fluid ascending
path consists of 60–100% serpentinite muds (Shipboard
Sci. Party 2002), the fluid-serpentinite interaction rather
than fluid-peridotite interaction (serpentinization) dominates the current fluid chemistry of the SCS formation
fluid. The highly alkaline porewater pH of ~ 12.5 likely prevents the development of subseafloor microbial community
due to the CO32−-dominated DIC speciation being disadvantageous for autotrophy (e.g., Schrenk et al. 2013)
whereas the known upper pH limit for chemoorganotrophic microorganisms is 12.4 (Takai et al. 2001). Ongoingreactions between the CH4 and SO4, as observed in thisstudy, would have been triggered by the very recent ODP
drilling and casing operations carried out in 2001 (Shipboard
Sci. Party 2002; Wheat et al. 2008).

subduction systems is a key factor for understanding
the (bio) geochemical cycle at forearc geofluid systems.
Constraint in keeping the co-existence of the (probably)slab-derived sulfate and abiotic CH4 in the SCS formation,
unexpected due to the strong thermodynamic drive for
their reaction, remains unidentified. Broadly speaking,abiotic organic synthesis associated with subseafloor
serpentinization has been considered to progress in manyof the natural serpentinization-associated geofluid systems
including the SCS, but has never been confirmed by multiple lines of solid evidences that serve to distinguish it fromorganic matter degradation and biological carbon fixation.
Future studies of the SCS and other serpentinite mud volcanoes in the Mariana forearc are needed to refine conceptual models of reactions within these interesting featuresthat link seafloor materials with the descending slab.

Conclusions
The present study investigated the chemical and microbiological characteristics of formation fluids from a serpentinite
mud volcano, the SCS. The gas chemistry characterized byabundant CH4 and low H2 is the first of its kind fromdeep-sea highly alkaline (serpentinization-associated) geofluid systems. The SCS formation fluid regime may also represent the first reported example on Earth where abiotic
processes are conspicuous with little biotic processes, despite
conditions with sufficient bioavailable energy potentials andtemperatures within the habitable range. However, there arestill many uncertainties associated with the hypothesis presented here. It is not sufficient to clearly distinguish the geochemical processes in each of the spatio-temporally different
stages of fluid-rock interactions beneath the SCS, such as
metamorphism associated with subduction, fluid released
from the plate boundary zone, extensive serpentinization at
the wedge mantle, fluid maturation and cooling in the ascending pathway, episodic mud volcanism, and mass
wasting of the SCS body. The behavior of sulfate inFunding
This work was partly supported by JSPS KAKENHI Grant Number 25701004 (SK).

Additional file
Additional file 1: Table S1. Results of chemical analyses of the CORK
fluid (XLSX 49 kb)
Abbreviations
CORK: Circulation Obviation Retrofit Kit; DPW: Deep porewater; IC: Ion
chromatography; ICP-AES: Inductively coupled plasma atomic emission
spectrometry; ICP-MS: Inductively coupled plasma mass spectrometry;
ODP: Ocean Drilling Program; PVC: Polyvinyl chloride; ROV: Remotely
operated vehicle; SCS: South Chamorro SeamountAcknowledgements
We thank the operation team of the ROV Hyper-Dolphin and the crews of R/V
Natsushima for their skillful support during cruises NT09-1, NT09-7, NT12-04, andNT12-23. We also thank Dr. Uta Konno for gas and isotope analyses. Drs. Kentaro
Nakamura, Takazo Shibuya, and Chong Chen are gratefully acknowledged for
useful and fruitful discussions. SK thanks Prof. Gretchen Bernasconi-Green forproviding space and opportunity to draft the manuscript as an academic visitorof ETH Zürich.

Availability of data and materials
All analytical results of this study are included in the main text and tables.
Authors’ contributions
SK, JM, CGW, and KT proposed the topic, conceived, and designed the study.
All authors joined cruises for sample collection. JM, YM, JSS, and CGW
analyzed the samples and helped in their interpretation. JSS and CGW
collaborated with the corresponding author (SK) in the construction ofmanuscript. All authors read and approved the final manuscript.
Competing interests
The authors declare that they have no competing interest.

Publisher’s NoteSpringer Nature remains neutral with regard to jurisdictional claims inpublished maps and institutional affiliations.
Author details
1
Department of Subsurface Geobiological Analysis and Research (D-SUGAR),
Japan Agency for Marine-Earth Science and Technology (JAMSTEC), 2-15
Natsushima-cho, Yokosuka 237-0061, Japan. 2Kochi Institute for Core Sample
Research, Japan Agency for Marine-Earth Science and Technology (JAMSTEC),
Nankoku, Kochi 783-8502, Japan. 3Department Marine Chemistry andKawagucci et al. Progress in Earth and Planetary Science(2018) 5:74

Geochemistry (MC&G), Woods Hole Oceanographic Institution (WHOI),
Woods Hole, MA 02543, USA. 4Global Undersea Research Unit, University of
Alaska Fairbanks, P. O. Box 475, Moss Landing, CA 95039, USA. 5Institute ofGeochemistry and Petrology, ETH Zürich, Zürich, Switzerland.
Received: 9 May 2018 Accepted: 19 October 2018

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