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Marine Geology

journal homepage: www.elsevier.com/locate/margo

Tectonic, climatic, and diagenetic control of magnetic properties ofsediments from Kumano Basin, Nankai margin, southwestern Japan

Meinan Shia,b, Huaichun Wua,b,⁎, Andrew P. Robertsc, Shihong Zhanga, Xixi Zhaod,e, Haiyan Lia,Xin Sub, Tianshui Yanga, Liao Changf, Pengxiang Huc, Xiang Zhaoc, Hongqiang Wangg

a State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Beijing 100083, Chinab School of Ocean Sciences, China University of Geosciences, Beijing 100083, Chinac Research School of Earth Sciences, Australian National University, Canberra, ACT 2601, Australiad State Key Laboratory of Marine Geology, Tongji University, Shanghai 200092, Chinae Earth and Planetary Sciences Department and Institute of Geophysics and Planetary Physics, University of California, 1156 High Street, Santa Cruz, CA 95064, USAf School of Earth and Space Sciences, Peking University, Beijing 100083, Chinag Key Laboratory of Seismic Observation and Geophysical Imaging, Institute of Geophysics, China Earthquake Administration, Beijing 100081, China

A R T I C L E I N F O

Keywords:Magnetic mineralsGas hydrateAnaerobic oxidation of methaneForearcKumano BasinNankai Trough

A B S T R A C T

Integrated Ocean Drilling Program (IODP) Site C0002 was cored in the Kumano Basin, which is a forearc basin inthe Nankai margin of southwest Japan. The uppermost 1052.5 m of sediment in the accretionary prism, slope,and overlying forearc basin was recovered discontinuously during Expedition 315. Magnetic minerals from SiteC0002 were studied to understand sedimentary, tectonic, and diagenetic processes in this forearc environment.Four rock magnetic units (A, B, C, and D) are identified. In the forearc basin sediments of Unit A, high con-centrations of detrital magnetic iron oxides in surface sediments give way to progressive diagenetic dissolutionwith depth. Nevertheless, large titanomagnetite particles survived diagenetic dissolution along with magneticnanoinclusions within host silicates that preserve a relict magnetic signal that records climatic modulation ofsediment delivery to the Kumano Basin. The authigenic iron sulphides, ferrimagnetic greigite and paramagneticpyrite, formed in association with gas hydrates and anaerobic oxidation of methane in the forearc basin sedi-ments of Unit B. Methane venting from deeper within the slope sediments of Unit C and accretionary prismsediments of Unit D has given rise to formation of pyrite aggregates along sediment fractures and has contributedto further diagenetic depletion of magnetic signals in these units. Structural evolution of the Kumano Basin hascaused migration of the gas hydrate stability zone (GHSZ) through time so that the Site C0002 sediments recordboth the present and fossil GHSZ positions. Magnetic properties of Kumano Basin sediments provide insights intothe interplay between diagenetic, climatic, and tectonic processes that controlled the sedimentary evolution ofthis forearc basin.

1. Introduction

Forearc basins form between oceanic trenches and associatedmagmatic arcs along convergent tectonic margins. Stratigraphic se-quences in such basins record the history of convergence, includingcollision of continental lithospheric plates, and oceanic plates withcontinents. They also reflect interactions between sedimentation, tec-tonics, climate, and sea-level fluctuations (Moore et al., 2015). How-ever, it is difficult to study these basins systematically because theyhave dynamic geological environments and the sedimentary record inancient forearc basins tends not to be well preserved (Ramirez et al.,2015). The Kumano forearc basin is located in the Nankai margin of

southwest Japan, where the Philippine Sea Plate (PSP) is subductingbeneath the Eurasian Plate (Fig. 1a; Seno et al., 1993; Miyazaki andHeki, 2001). It has been the subject of recent attention as a source ofgreat thrust earthquakes and for commercial utilization of gas hydrates(Tobin and Kinoshita, 2006; Miyakawa et al., 2014; Wiersberg et al.,2015). As part of the Integrated Ocean Drilling Program (IODP) NankaiTrough Seismogenic Zone Experiment (NanTroSEIZE), Sites C0009 andC0002 were drilled in the Kumano forearc basin (Expedition 315Scientists, 2009; Expedition 319 Scientists, 2010; Strasser et al., 2014).The tectonic and sedimentary history of the Kumano Basin can now beassessed in detail by combining 3D and 2D seismic reflection data withdrill hole lithological, geochemical, and geophysical data (Park et al.,

http://dx.doi.org/10.1016/j.margeo.2017.07.006Received 21 March 2017; Received in revised form 3 July 2017; Accepted 6 July 2017

⁎ Corresponding author at: School of Ocean Sciences, China University of Geosciences, Beijing 100083, China.E-mail address: whcgeo@cugb.edu.cn (H. Wu).

Marine Geology 391 (2017) 1–12

Available online 13 July 20170025-3227/ © 2017 Elsevier B.V. All rights reserved.

MARK

2002; Moore et al., 2007, 2015; Strasser et al., 2009; Kimura et al.,2011; Underwood and Moore, 2012; Usman et al., 2014; Buchs et al.,2015; Ramirez et al., 2015).

Magnetic minerals in marine sediments originate mainly from det-rital input, biomineralization, or post-depositional alteration. Thecomposition and grain size of magnetic minerals are sensitive to en-vironmental changes. Therefore, magnetic properties can help to tracksediment sources and transportation pathways and to characterize de-positional regimes and geochemical environments (e.g., Thompson andOldfield, 1986; Hilgenfeldt, 2000; Rowan et al., 2009; Just et al., 2012;Liu et al., 2012; Roberts et al., 2013). For example, dissolution of pri-mary magnetic minerals and authigenic growth of secondary magneticiron sulphides are usually used to detect sulphidic diagenetic conditionsor the presence of gas hydrates (Housen and Musgrave, 1996; Musgraveet al., 2006; Enkin et al., 2007; Larrasoaña et al., 2007; Rowan et al.,2009; Kars and Kodama, 2015a, 2015b; Roberts, 2015). Recently, Karsand Kodama (2015a, 2015b) studied the magnetic characteristics of gashydrate-bearing sediments at IODP Site C0008 from the megasplay faultzone of Nankai Trough and revealed the importance of diageneticprocesses. However, knowledge of the magnetic properties of marinesediments in forearc basins and their relationship to tectonic evolutionremains poor. Previous magnetic studies of sites C0002 and C0009 re-veal the presence of magnetite and iron sulphides (Expedition 315Scientists, 2009; Expedition 319 Scientists, 2010; Kanamatsu et al.,2012). The sampling resolution of these preliminary studies, however,was low, and they did not focus on diagenetic processes or their en-vironmental implications.

We present detailed magnetic results and electron microscope ob-servations for 374 discrete samples from Site C0002. Our objectives areto: 1) obtain a detailed magnetic parameter sequence for a verticalsection through the Kumano Basin, and to 2) reveal relationships

between magnetic mineral assemblages, sedimentation, tectonic evo-lution, climatic modulation, and diagenesis.

2. Geological setting and sampled material

Nankai Trough is a convergent margin, where the Philippine SeaPlate is subducting to the northwest beneath the Eurasian Plate at a rateof ~4 cm/yr (Fig. 1a; Seno et al., 1993; Miyazaki and Heki, 2001). Theconvergence direction is approximately normal to the trench and the KiiPeninsula, and sediments of Shikoku Basin are actively accreting at thedeformation front (Park et al., 2002). Kumano Basin sits between theupper trench slope and a megasplay fault associated with a forearc highthat marks the transition in the Nankai margin from a downslope to animbricate thrust zone. The basin basement is a Miocene–Pliocene ac-cretionary prism that is overlain by lower Pliocene trench-slope de-posits (Expedition 315 Scientists, 2009).

Uplift along a regional out-of-sequence thrust (megasplay) faultbegan at around 1.95 Ma (Strasser et al., 2009) and is thought to havecreated accommodation space for thick forearc-basin deposits near thedistal edge of the basin on the Kumano transect (Fig. 1b; Underwoodand Moore, 2012). Motion along a new, landward splay fault branchcaused tilting of the lower forearc basin between ~1.2 and 0.9 Ma(Gulick et al., 2010); continued splay fault slip caused further landwardtilting of the sediment sequences (Gulick et al., 2010; Moore et al.,2015). A gas hydrate-bearing zone was identified indirectly in the in-terval from 218 to 401 m below seafloor (mbsf) above a bottom-simu-lating reflector (BSR) identified from electrical resistivity and seismic P-wave velocity measurements (Expedition 314 Scientists, 2009). Massivemethane release is considered to have occurred along the seaward edgeof the Kumano forearc basin at 0.05 Ma (Bangs et al., 2010).

Site C0002 (Fig. 1a, b) is located on the southern margin of Kumano

Dep

th (

km)

10

8

6

4

2

5 km

VE = 1.5X

Subducting Philippine Sea Plate

Older accretionary prism Accretionary prism

Shikoku Basin deposits

“notch”C0004C0002

Kumano Basin

kmTokai

~ 4 cm/yr

0 50

33°

34°

35° N

Philippine Sea Plate(Shikoku Basin)

KiiPeninsula

40°

N

30°

130 E° 140° 150°

Ise Bay

Kumano Basin

Site C0002Site C0003-5,C0008

Site C0006, C0007

Site C0011

Site C0012

Site C0009

135° E 136° 137° 138° 139°

(a)

(b) 7000C6000C1000C9000CC0008

Fig. 1. (a) Location of IODP Expedition 315 Site C0002 andother drill sites from the Nankai margin, offshore of SW Japan.(b) Spliced composite profile of a representative 3D seismicreflection depth section with positions of IODP drill Site C0002and several other drill sites (after Underwood et al., 2009).

M. Shi et al. Marine Geology 391 (2017) 1–12

2

Basin. The uppermost 1052.5 m of the sequence recovered during IODPExpedition 315 is divided into four lithological units according to welllogging data and lithological changes (Fig. 2). The upper forearc basinfacies (Unit I; 0–135.85 m CSF (core depth below seafloor)) consists ofQuaternary hemipelagic silty mud with multiple interbedded sandy andvolcanic ash layers. The lower forearc basin facies (Unit II;135.85–834.00 m CSF) is characterized by silty hemipelagic mud withoccasional interbedded sandy and volcanic ash layers. An interval inUnit II (204–475 m CSF) was not recovered until IODP Expedition 338(Strasser et al., 2014). Hence in this study, no samples come from thisinterval. The Pliocene slope facies of Unit III (834.00–921.70 m CSF) ischaracterized by a condensed mudstone section. The Miocene accre-tionary prism of Unit IV (921.7–1052.5 m CSF) consists of hemipelagicmudstone with some siltstone and sandstone. The accretionary prismhas been subjected to deformation, as evidenced by steepened bedding,faults, breccia, and shear zones (Expedition 315 Scientists, 2009). De-positional ages at Site C0002 were determined based on nannofossil andforaminifer ages (Fig. 2; Expedition 315 Scientists, 2009). The magneticpolarity stratigraphy is ambiguous (Strasser et al., 2014), so it was notused here for age model construction.

3. Methods

In this study, 374 discrete specimens were collected from holesC0002B and C0002D at ~1 m sampling intervals with non-magnetic2 × 2 × 2 cm plastic cubes. The sample naming convention was: hole-core-section-interval-depth, where the interval is in centimetres belowthe top of the section; e.g., C0002B-20R-1W, 70.0–72.0 cm, 647.71 m

CSF. The dominant lithology of the samples is grey mud (silty clay toclayey silt). Low-frequency (976 Hz) magnetic susceptibility (χ, mass-specific) was measured using a MFK1-FA Kappabridge magnetic sus-ceptibility meter at room temperature. An anhysteretic remanentmagnetization (ARM) was acquired by applying a peak alternating field(AF) of 100 mT and a direct current (DC) bias field of 0.05 mT using aD-2000 AF demagnetizer. ARM values are expressed as the suscept-ibility of ARM (χARM = ARM/DC bias field). Saturation isothermal re-manent magnetization (SIRM) acquired in a field of 2.0 T was appliedusing an ASC IM-10-30 impulse magnetizer. Remanent magnetizationmeasurements were made with a JR-6A spinner magnetometer. An IRMimparted at 300 mT (IRM-300mT) in the opposite direction to the SIRMwas measured to calculate the S-ratio = (− IRM-300mT/SIRM).Temperature-dependent susceptibility (χ–T) was measured betweenroom temperature and 700 °C in an argon atmosphere using a KLY-4SKappabridge system with a CS-3 high-temperature apparatus. Thesemagnetic analyses were performed at the Paleomagnetism andEnvironmental Magnetism Laboratory at China University ofGeosciences, Beijing.

Hysteresis loop, first-order reversal curve (FORC) and IRM acqui-sition measurements were made on representative samples at roomtemperature using a MicroMag Model 3900 vibrating sample magnet-ometer (VSM) at the Australian National University (ANU), Canberra,Australia. Hysteresis loops were measured between± 500 mT with afield step of 3 mT and 500 ms averaging time. Hysteresis parameters,including saturation magnetization (Ms), saturation remanent magne-tization (Mrs), and coercive force (Bc) were determined from the hys-teresis loops. FORCs were measured using the irregular FORC protocol,

0 1 2 3 4 5 6

Biostratigraphic age model (Ma)

QuaternaryenecoiMenecoilP

atelyeta lrael middle

NN11b

NN16NN17

NN19

NN21

NN18

NN15-12

Minimum age

Minimum age

Maximum age

Nannofossil zone

Unit II

Unit III

Unit IV

Unit I

ForaminiferNannofossil Depth range of event

50-300 m/Myr

400-800 m/Myr

18-30 m/Myr

Lithology

Dep

th C

SF (

m)

Silty claySand Volcanic ash BrecciaSilty, sandy silt

0

200

400

600

800

1000

B D H J K L

Exp 3

15

Recovery Hole

Fig. 2. Lithology and age model for IODP Site C0002;CSF = core depth below seafloor (from Expedition 315Scientists, 2009; Strasser et al., 2014). Black = core recovered.Grey = interval cored. The age model is based on biostrati-graphy.

M. Shi et al. Marine Geology 391 (2017) 1–12

3

in which irregularly spaced field steps are defined based on the majorhysteresis loop properties (Zhao et al., 2015). Data were processedusing the iFORC software (Zhao et al., 2015) with smoothing factor (SF)of 3. IRM acquisition curves were obtained up to maximum fields of 1 T.IRM acquisition curves were decomposed into magnetic coercivitycomponents using the fitting protocol of Kruiver et al. (2001). Re-manent coercive force (Bcr) was determined from IRM acquisitioncurves.

Low-temperature magnetic measurements were performed with aQuantum Design Magnetic Properties Measurement System (MPMS XL-5) in the Paleomagnetism and Geochronology Laboratory, Institute ofGeology and Geophysics, Chinese Academy of Sciences, Beijing. SIRMwas imparted in a 2.5 T field at 15 K. Zero-field-cooled (ZFC) and field-cooled (FC) curves were obtained during cooling from 300 to 15 K inzero field and in a 2.5 T field, respectively, and then by measuring theremanence in zero field during warming to 300 K.

Magnetic mineral extraction was performed following the procedureof Chang et al. (2012). Resin-impregnated bulk samples and magneticparticles were investigated using a Zeiss SUPPA 55 field emissionscanning electron microscope (FESEM) with electron diffraction spec-trometer (EDS) at the FESEM Laboratory, China University of Geos-ciences, Beijing. Samples were attached to a SEM stub with conductiveadhesive and were coated with platinum. Secondary electron (SE) andbackscattered electron (BSE) imaging modes were used. Transmissionelectron microscope (TEM) observations of magnetic extracts werecarried out at the Centre for Advanced Microscopy, ANU, with a PhilipsCM300 TEM operated at 300 kV with an EDAS Phoenix retractable X-ray detector.

4. Results

4.1. Down-core magnetic parameter variations

Down-core profiles of magnetic parameters are shown in Fig. 3.Rock magnetic units are named Units A, B, C, and D, with boundaries at105.00, 834.00, and 921.70 m CSF, respectively. Shipboard κ andnatural remanent magnetization (NRM), measured before and after AFdemagnetization at 20 mT during IODP Expeditions 315 and 338(Expedition 315 Scientists, 2009; Strasser et al., 2014), have similarvariations, which suggest that their variations are controlled mainly bymagnetic mineral concentration. NRM after demagnetization at 20 mTshows that the influence of a coring-induced overprint has been re-moved successfully. Magnetic concentration-dependent parameters χ,ARM, and SIRM also have the same down-hole trends, with low valuesin Units A, C, and D.

In Unit A, magnetic parameters decrease progressively down-coreand then increase again and peak at ~60 m CSF. In Unit B, theseparameters have high values between 110 and 650 m CSF, and thenfluctuate significantly between 650 and 820 m CSF. In Units C and D,magnetic parameter values are relatively stable. SIRM in Unit D is onaverage lower than in other units. χARM/χ has high values in Unit B,while ARM/SIRM has low values. SIRM/χ is also an indicator of mag-netic mineral grain size, where high SIRM/χ values are often used toindicate the presence of stable single domain (SD) magnetic iron sul-phides (e.g., Snowball and Thompson, 1988; Snowball, 1991; Roberts,1995; Larrasoaña et al., 2007; Roberts et al., 2011). SIRM/χ is gen-erally< 20 kA/m in Units A, C, and D. Samples from Unit B havevariable SIRM/χ values, mostly> 50 kA/m. The S-ratio indicates therelative concentration of low- to high-coercivity minerals. S-ratioscentre on 0.85 in Units A and C and are mainly larger than 0.95 in UnitsB and D, which suggests the dominance of low coercivity minerals inthese units and a greater concentration of high coercivity minerals inUnits A and C.

4.2. Temperature-dependent magnetic properties

Temperature-dependent magnetic measurements can be diagnosticof magnetic mineralogy. Magnetic susceptibility drops to zero at~550–580 °C in heating curves for all samples in this study, which isthe Curie temperature of (titano-) magnetite. Most χ-T curves increaseat ~300 °C, reach a maximum at ~500 °C, and gradually decrease to580 °C (Fig. 4). This suggests the presence of iron sulphides (e.g.,mainly pyrite) that decompose and convert to magnetite during heating(e.g., Passier et al., 2001). For samples from Units C and D, coolingcurves increase strongly between 600 °C and 300 °C with respect to theheating curves. This suggests that magnetite formed during heating.The Verwey transition at ~120 K due to magnetite (Verwey, 1939;Muxworthy and McClelland, 2000) is recognizable for all studiedsamples in both ZFC and FC curves. An unblocking temperature in thevicinity of 30–40 K is most likely caused by superparamagnetic particles(Moskowitz et al., 1998).

4.3. Hysteresis properties and FORC diagrams

Hysteresis loops and FORC diagrams help to discern the magneticdomain state (i.e., grain size) of magnetic minerals. Pink and bluehysteresis loops (Fig. 4) are normalized to Ms (at 500 mT) before andafter paramagnetic slope correction, respectively. The magnetizationwas saturated below 70% of the maximum applied field, so the high-field slope above this value was used for correction. Corrected hyster-esis loops from Units A, C, and D are narrow with Mrs/Ms of~0.20–0.32 and Bc of ~13–25 mT. Corrected hysteresis loops fromUnit B generally have relatively square shapes with high Mrs/Ms of0.33–0.54 and Bc of 23–44 mT. Hysteresis loops from the base of UnitsB to D have a dominant paramagnetic contribution, and low Bc.

A FORC diagram for a sample from uppermost Unit A (Fig. 4) pro-vides evidence of a broad coercivity distribution between 0 and 80 mTand a peak Bc of ~20 mT. The vertical distribution along the Bu axis isbroad. This pattern is characteristic of pseudo-single domain (PSD)particles (Roberts et al., 2000, 2014; Muxworthy and Dunlop, 2002). AFORC diagram for a sample from the middle of Unit A has an elongatedcoercivity distribution between 0 and 200 mT, which is indicative of ahigh coercivity mineral (e.g., hematite) (Roberts et al., 2006). FORCdiagrams from Unit B (Fig. 4) generally have concentric contours withlarge vertical spread, which represent a strongly interacting SD com-ponent with peak Bc values at ~40 mT. This is characteristic of greigite(Roberts et al., 2006, 2011). FORC diagrams for Units C and D have a Bc

peak at ~20 mT, elongated Bc distributions between 0 and 160 mT, andnegligible magnetostatic interactions, which are representative of fine-grained titanomagnetite with a small high coercivity mineral content.

4.4. IRM acquisition

IRM acquisition curves for the studied samples (Fig. 4) can be fitted(Kruiver et al., 2001) with 1 to 3 components. A soft component 1(cyan) with B1/2 down to 1 mT is interpreted to represent an asym-metric base function with no physical meaning (Egli, 2004). Component2 (blue) is interpreted as magnetite, with a 15–50 mT mean coercivity.The mean coercivity of component 3 (yellow) is ~60–80 mT, which isin the range expected for greigite. The mean coercivity of the fourthcomponent (red) is> 100 mT, and is interpreted to represent hematite.

4.5. Electron microscope observations

Fragmented titanomagnetites (Fig. 5a) are the most frequently ob-served magnetic particles in magnetic mineral extracts from Units A toC. Titanomagnetite occurs in grain size ranges from tens of nanometres(nm) to tens of micrometres (μm), which spans the superparamagneticto multi-domain range for magnetite (Dunlop and Özdemir, 1997). Acommonly observed magnetic particle type in Unit A is euhedral

M. Shi et al. Marine Geology 391 (2017) 1–12

4

titanomagnetite with typical octahedral spinel-type shapes (Fig. 5c, d).The observed titanomagnetites have sizes of ~1–50 μm, which is con-sistent with the dominantly PSD magnetic properties in Unit A (Fig. 4).Clusters of fine-grained iron sulphide nanoparticles (Fig. 5f) are iden-tified in Unit A. Greigite is detected in Unit B, and pyrite framboids arefrequently found in lower Unit B (Fig. 5h–k). Fragmented titano-magnetite particles (Fig. 5b) with pits or crevices on their surfaces arecommon in Unit C. A small amount of greigite and pyrite (Fig. 5l) isfound. Numerous pyrite aggregates (Fig. 5n, o) in Unit D occur alongfractures. Some pyrite framboids (Fig. 5m) have sizes> 10 μm withnegligible overgrowths. In addition, TEM observations indicate thepresence of nanoparticle magnetite inclusions within host silicate grains(Fig. 5p).

5. Discussion

As outlined below, the magnetic properties of the studied sedimentsat IODP Site C0002 are controlled by a range of factors, includingtectonic and climatic forcing of erosion, transportation, and depositionof detrital particles, and by diagenesis. In the following discussion, wedescribe the range of documented processes from younger to older ra-ther than with time (older to younger) because recent processes providean understanding of those documented lower in the hole.

5.1. Diagenetic dissolution of detrital magnetic minerals

As is observed commonly in rapidly deposited marine sedimentswith moderate to high organic carbon contents (which averages 0.52 wt% at Site C0002; Fig. 3; Expedition 315 Scientists, 2009), initially high

110

100

1000

(10-5

SI)

1000900

800

700

600

500

400

300

200

100

DepthCSF(m)

Period/Epoch

Lithology

Rockmagneticunit

Quaternary Plio. Mio.

A B C D

10m

3 kg-1

)

110

100

AR

M(1

0-6A

m2 k

g-1)

100

1 000

1000

0

SIR

M10

-6A

m2 k

g-1 110

AR

M/

10

AR

M/S

IRM

10-3

110

100

SIR

M/

(kA

/m)0.

70 .

80.

91

S-ra

tio

01 0

2030

SO4

050

0 00

Met

hane

(pp m

v)

00.

81.

62 .

4T

OC

(wt%

)

00.

81.

62.

4

TS

(wt %

)

Gas

hydr

ate

zone

503

505

0.01

11 0

0N

atur

alre

man

ent m

agne

tizat

ion

AF

20m

T d

emag

netiz

edin

tens

ity

(10

-3A

/m)

Relativelocations ofholes

D

K L B

Ship

boar

dFig. 3. Down-core variations of magnetic parameters for SiteC0002. From left to right: shipboard magnetic measurementsrecovered during IODP Expedition 315 and 338 (κ, NRM in-tensity before (blue) and after (dark red) AF demagnetization at20 mT) (Expedition 315 Scientists, 2009; Strasser et al., 2014),magnetic concentration-dependent parameters (χ, ARM, andSIRM), magnetic granulometric parameters (χARM/χ, ARM/SIRM, and SIRM/χ), magnetic composition-dependent para-meter (S-ratio), and pore water geochemical variations, in-cluding sulphate and methane, and total organic carbon (TOC)and total sulphur (TS) (Expedition 315 Scientists, 2009; Strasseret al., 2014). The scale for ARM/SIRM is inverted. (For inter-pretation of the references to colour in this figure legend, thereader is referred to the web version of this article.)

M. Shi et al. Marine Geology 391 (2017) 1–12

5

0 40 80 120Bc (mT)

-80

-40

0

40

80

Bu

(mT

)

-1

-0.5

0

0.5

1

SF = 3

Unit A

Unit B

Unit C

Unit D

0 1 2 3Applied field (Log10 mT)

0

0.4

0.8

1.2

Nor

mal

ised

IRM

Gra

dien

t

-400 -200 0 200 400Field (mT)

-0.8

-0.4

0

0.4

0.8

Mrs

/Ms

Bc = 13.1 mT

Mrs/Ms = 0.204

6

8

Mrs

(0.1

A/m

)0 100 200 300

Temperature (K)

-120

-80

-40

0

/mK

ZFCFC

0 200 400 600Temperature (oC)

0

0.3

0.6

0.9

/o

C0002D-1H-1W, 53.0-55.0 cm,0.54 m CSF

C0002D-6H-8W, 74.0-76.0 cm,52.10 m CSF

0 200 400 600Temperature (oC)

0

0.3

0.6

0.9

/o

C0002D-14H-7 W, 61.0-63.0 cm,126.66 m CSF

0 200 400 600Temperature (oC)

0

0.3

0.6

0.9

/o

C0002B-1R-3 W, 70.0-72.0 cm,477.45 m CSF

0 200 400 600Temperature (oC)

0

0.3

0.6

0.9

/o

C0002B-31R-2 W, 53.0-55.0 cm,752.95 m CSF

0 200 400 600Temperature (oC)

0

0.3

0.6

0.9

/o

C0002B-45R-2 W, 56.0-58.0 cm,883.485 m CSF

0 200 400 600Temperature (oC)

0

0.3

0.6

0.9

/o

C0002B-55R-CC W, 7.0--9.0 cm,978.44 m CSF

0 200 400 600Temperature (oC)

0

0.3

0.6

0.9

/o

0

2

4

6

Mrs

(0.1

A/m

)

0 100 200 300Temperature (K)

-100

-80

-60

-40

-20

0

/mK

ZFCFC

0

2

4

6

Mrs

(0.1

A/m

)

0 100 200 300Temperature (K)

-60

-40

-20

0

/mK

ZFCFC

0

4

8

12

16

Mrs

(0.1

A/m

)

0 100 200 300Temperature (K)

-30

-20

-10

0

/mK

ZFCFC

4

8

12

16

Mrs

(0.1

A/m

)

0 100 200 300Temperature (K)

-120

-80

-40

0

/mK

ZFCFC

2

4

6

Mrs

(0.1

A/m

)

0 100 200 300Temperature (K)

-40

-30

-20

-10

0

/mK

ZFCFC

0

4

8

12

Mrs

(0.1

A/m

)

0 100 200 300Temperature (K)

-60

-40

-20

0

/mK

ZFCFC

-400 -200 0 200 400Field (mT)

-0.8

-0.4

0

0.4

0.8

Mrs

/Ms

Bc = 25.3 mT

Mrs/Ms = 0.32

-400 -200 0 200 400Field (mT)

-0.8

-0.4

0

0.4

0.8

Mrs

/Ms

Bc = 43.6 mT

Mrs/Ms = 0.49

-400 -200 0 200 400Field (mT)

-0.8

-0.4

0

0.4

0.8

Mrs

/Ms

Bc = 41.1 mT

Mrs/Ms = 0.54

-400 -200 0 200 400Field (mT)

-0.8

-0.4

0

0.4

0.8

Mrs

/Ms

Bc = 23.4 mT

Mrs/Ms = 0.33

-400 -200 0 200 400Field (mT)

-0.8

-0.4

0

0.4

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magnetic mineral concentrations in surficial sediments give way to agradual decrease in magnetization with depth. In Unit A of HoleC0002D, the minimum magnetization occurs when pore water sulphatevalues decrease to nearly 0 μM at depths of ~20 m CSF (Fig. 3;Expedition 315 Scientists, 2009). This is a typical signature associatedwith magnetite dissolution in sulphate-reducing sediments (Karlin andLevi, 1983; Canfield and Berner, 1987; Karlin, 1990a, 1990b; Channelland Hawthorne, 1990; Roberts, 2015). Once the detrital magnetic mi-neral assemblage has undergone dissolution, underlying sedimentsusually have extremely weak magnetizations. This is not observed inHole C0002D (Fig. 3). Large-scale fluctuations of magnetic concentra-tion-dependent parameters are superimposed on the overall trend ofdeclining magnetizations from the core surface to depths of ~20 m CSFand in the underlying diagenetically reduced sediments (Fig. 3). Thissuggests that despite the influence of sulphate-reducing conditions andubiquitous accompanying magnetite dissolution (Canfield and Berner,1987; Roberts, 2015), significant detrital magnetic mineral concentra-tions have survived diagenetic dissolution. Two lines of evidence ex-plain this observation. First, large titanomagnetite particles are pre-served in the uppermost part of Unit A (Fig. 5c), and within the sulphate

reduction zone at 20.22 m CSF, although the latter particles have pittedsurfaces or rounded edges probably due to corrosion by sulphidic porefluids (Fig. 5a, d). Thus, while authigenic iron sulphides have formed atthe expense of detrital iron-bearing particles throughout the upper partof Hole C0002D (Fig. 5f, g), it appears that the large particle sizes andthe relatively lower reactivity of titanomagnetite grains with respect tostoichiometric magnetite (Poulton et al., 2004; Roberts, 2015) meansthat they have not undergone complete dissolution. Large titano-magnetite particles also occur in Units B and C (Fig. 5b, e), whichsupports this conclusion. Second, titanomagnetite occurs as nano-particle inclusions within host silicates in TEM images (Fig. 5p). Changet al. (2016a) reported similar observations from the Japan margin tothe north (offshore of Tokyo) and demonstrated that (titano-) magnetiteoccurs commonly as inclusions within silicate host particles that protectdetrital magnetic minerals from dissolution. Magnetic nanoinclusionswithin host silicates are common in igneous rocks and should also bepreserved widely within detrital silicates in sediments (Chang et al.,2016a, 2016b). Thus, the two documented types of magnetic mineraloccurrences are responsible for persistent detrital magnetic signalswithin sediments that have undergone sulphidic diagenesis.

Fig. 4. Mineral magnetic results for typical samples from magnetic Units A to D. Magnetic data in each row are from the same sample. Column 1: high-temperature magnetic susceptibility(χ-T) curves. Red and blue lines represent heating and cooling curves, respectively. Column 2: low-temperature magnetization curves (ZFC and FC; blue) and their derivatives (ΔM/ΔT;pink). The derivative curves contain an inflection at the Verwey transition temperature (~120 K), which indicates the presence of magnetite in all samples. Column 3: hysteresis loops.Pink and blue loops are normalized to Ms values (at 500 mT) before and after paramagnetic slope correction (corrected from the slope above 70% of the maximum applied field where themagnetization is saturated), respectively. Column 4: first-order reversal curve (FORC) diagrams (SF = 3). Column 5: isothermal remanent magnetization (IRM) component analyses. Greysquares represent data points that define the gradient of IRM acquisition curves. Cyan, yellow, blue, and red represent the identified coercivity components (see text for discussion). (Forinterpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Energy (keV)

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nts

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(k) C0002B-31R-2W, 53.0-55.0 cm, 752.95 m CSF (Unit B)

(l) C0002B-45R-2W, 56.0-58.0 cm, 883.48 m CSF (Unit C)

(n) C0002B-61R-4W, 49.0-51.0 cm, 1021.51 m CSF (Unit D )

(m) C0002B-55R-CCW, 7.0-9.0 cm, 978.44 m CSF (Unit D)

(a) C0002D-3H-5W, 90.0-92.0 cm, 20.22 m CSF (Unit A)

(b) C0002B-45R-2W, 56.0-58.0 cm, 883.48 m CSF (Unit C)

(d) C0002D-3H-5W, 90.0-92.0 cm, 20.22 m CSF (Unit A)

(c) C0002D-1H-1W, 53.0-55.0 cm, 0.54 m CSF (Unit A)

(e) C0002B-1R-3W, 70.0-72.0 cm, 477.45 m CSF (Unit B)

(f) C0002D-1H-1W, 53.0-55.0 cm, 0.54 m CSF (Unit A)

(g) C0002D-3H-5W, 90.0-92.0 cm, 20.22 m CSF (Unit A)

(h) C0002D-18H-4W, 18.0-20.0 cm, 202.22 m CSF (Unit B)

(i) C0002B-10R-1W, 68.0-70.0 cm, 552.69 m CSF (Unit B)

(j) C0002B-10R-1W, 68.0-70.0 cm, 552.69 m CSF (Unit B)

(o) C0002B-61R-4W, 49.0-51.0 cm, 1021.51 m CSF (Unit D )

Py

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SE SE SE SE SE

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BSE SE BSE RLM SE

RLM

TEM SE

Inclusion

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Fig. 5. Electron microscope images of representative samples from the four magnetic units (A–D) of Site C0002: (a–b) fragmented titanomagnetite particles, (c–e) euhedral titano-magnetite, (f–o) iron sulphide aggregates, and (p) titanomagnetite nanoparticles within a silicate particle. Red and purple squares in (p) indicate EDS analysis spots for titanomagnetitenanoparticles, as presented in (q, r). Cyan and blue squares indicate EDS analysis spots for the host particle, as presented in (s, t). The inclusions contain Fe and O, so are interpreted torepresent magnetite. The host mineral only contains Si and O. Images in (a) to (o) are from FESEM observations, and images in (p) are from TEM observations with the spectra in (q) to (t)representing associated EDS analyses. Ti-Mg: titanomagnetite, Py: pyrite, G: greigite, SE: secondary electron image, BSE: backscattered electron image, TEM: transmission electronmicroscope image, RLM: reflected light micrograph. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

M. Shi et al. Marine Geology 391 (2017) 1–12

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5.2. Climatic modulation of detrital magnetic mineral supply to the KumanoBasin

In addition to reporting that the magnetic signal from sulphidicsediments is carried by magnetic nanoinclusions within silicate hostparticles, Chang et al. (2016a) found that delivery of such erosionaldetritus from Japan to its Pacific margin and supply of organic carbonto the seafloor was modulated climatically by millennial scale AsianMonsoon fluctuations through the last glacial cycle. These variationsdrove a non-steady state diagenetic response that was recorded by therelict magnetic mineral assemblages preserved as magnetic nanoinclu-sions within silicate host particles. Variations in sea level and changesin sediment supply routes from different sources, and tectonic changes,are all further possible causes of sedimentary variations in such settings.The possibility of climatic modulation of sedimentation is explored inFig. 6.

We restrict our analysis to the most recent time interval (0–800 ka)

over which a reasonable age model can be constructed from depth-based δ18O variations in planktonic foraminifera (Fig. 6f) (Matsuzakiet al., 2015). Our sampling resolution is relatively low, so we use κ(Expedition 315 Scientists, 2009) as a high-resolution proxy for mag-netic mineral content variations in the upper part of Hole C0002D,which has a positive relationship with χ and all other concentration-dependent magnetic parameters (Fig. 3). The κ data were filtered usinga Taner filter (Taner, 2000), and segmented precession cycles wereextracted from the filtered κ record (Fig. 6c, Fig. S1). A sedimentaryhiatus was reported in Hole C0002D at around 35–47 m CSF(Expedition 315 Scientists, 2009; Matsuzaki et al., 2015), so the recordsabove and below this interval were filtered separately. By comparingthe depth-based δ18O record from Hole C0002D (Fig. 6f) with the age-based global benthic δ18O stack (Fig. 6g; Lisiecki and Raymo, 2005) anda representative high-resolution composite Asian monsoon δ18O recordfrom Chinese speleothems (Fig. 6h; Cheng et al., 2016), higher mag-netic mineral contents (Fig. 6b) and generally finer magnetic minerals

0.8

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Fig. 6. (a) κ (Expedition 315 Scientists, 2009), (b) χ, (c) seg-mented precession cycles extracted from the κ series using Taner(2000) filters (Fig. S1), (d) χARM/χ, (e) S-ratio, (f) δ18O varia-tions in the planktonic foraminifer G. inflata (Matsuzaki et al.,2015), (g) the LR04 global benthic δ18O stack of Lisiecki andRaymo (2005), (h) the composite Asian monsoon δ18O spe-leothem record of Cheng et al. (2016), and (i) changes in orbitaleccentricity (red) and precession (blue) from Laskar et al. (2004)for the last 800 ka. The legend for lithology is the same as inFig. 2. (For interpretation of the references to colour in thisfigure legend, the reader is referred to the web version of thisarticle.)

M. Shi et al. Marine Geology 391 (2017) 1–12

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(Fig. 6d) are apparently preserved in sediments that were depositedduring drier summer monsoon periods (Fig. 6h; grey correlation lines inFig. 6). The age model is approximate, which precludes a 1:1 matchbetween κ peaks and drier monsoon intervals. In three intervals, low S-ratio values appear to correspond to drier periods (pink correlationlines in Fig. 6). The occurrence of higher coercivity magnetic mineralsin these intervals is probably related to eolian dust inputs due to driermonsoon conditions. Apart from this, it is not clear how Asian monsoonvariations have influenced magnetic mineral inputs to the KumanoBasin. This is particularly the case when a more likely relationshipwould be associated with wetter monsoon conditions that presumablyrelate to stronger run-off and erosion from Japan. In a tectonicallyactive margin, erosion of siliciclastic sediment from Japan would beexpected to dominate eolian inputs, so it is difficult to imagine a con-sistently detectable eolian contribution to sedimentation associatedwith drier monsoon conditions in this setting.

We suggest that the apparent link between drier Asian monsoonsand higher magnetic mineral inputs in the Kumano basin (Fig. 6) wasrelated to climate variations as follows. The most obvious links betweenκ peaks in Hole C0002D and the monsoon index of Cheng et al. (2016)occur during peak interglacial periods (grey correlation lines in Fig. 6).Interglacials correspond to sea level highstands (e.g., Rohling et al.,2014), so it is possible that sediment transportation pathways changedbetween glacial and interglacial periods. In such a scenario, sedimentstored on continental shelves during sea level lowstands will be redis-tributed when sea level rises and mobilizes sediment into adjacentmarine basins. Alternatively, warmer sea surface temperatures in in-terglacial periods make intense subtropical storms more likely. Suchsubtropical storm events enhance terrestrial erosion, which can result insignificant offshore sedimentation (e.g., Horng and Roberts, 2006). Ineither scenario, climatically modulated influxes of terrigenous silici-clastic detritus, which will contain silicates with magnetic nanoinclu-sions, will lead to higher magnetic mineral contents and finer magneticgrain sizes that will not be affected strongly by the prevailing sulphidicdiagenetic conditions that have affected these sediments since burial.

5.3. Diagenetic iron sulphide formation and tectonic forcing of gas hydratestability zones

Iron sulphides are found throughout Unit B. These minerals form inassociation with microbially mediated bacterial sulphate reduction(Roberts, 2015). During early diagenesis of continental margin en-vironments, iron sulphide precipitation is also commonly associatedwith hydrogen sulphide formation via anaerobic oxidation of methane(AOM) (Kasten et al., 1998; Neretin et al., 2004; Liu et al., 2004;Garming et al., 2005; Riedinger et al., 2005; Horng and Chen, 2006;Rowan and Roberts, 2006; Larrasoaña et al., 2007; März et al., 2008;Rowan et al., 2009; Roberts, 2015). The AOM reaction zone occursacross the sulphate-methane transition (SMT). The H2S produced reactswith iron-bearing minerals to cause dissolution of detrital magneticminerals and formation of end-member pyrite through the intermediateprecursor greigite. The depth at which the SMT occurs within sedimentsis mainly a function of organic matter content and reactivity, and se-dimentation rate (Roberts, 2015). Greigite and pyrrhotite can also formbelow the SMT as a result of microbial activity enhanced by methanehydrates or gas venting, which suggests that diagenetic greigite andpyrrhotite can be useful for detecting gas hydrate systems (Housen andMusgrave, 1996; Larrasoaña et al., 2007; Dewangan et al., 2013; Karsand Kodama, 2015a, 2015b; Roberts, 2015). Kars and Kodama (2015a,2015b) analyzed magnetic minerals in sediments from Site C0008 tothe south-east of Site C0002, and proposed that layers containinggreigite and/or pyrrhotite correspond to the presence of (fossil) gashydrate horizons. Release of H2S below the base of the gas hydrate layerallows conversion of greigite to pyrite. Iron sulphide formation is alsorelated to tectonically driven methane migration. The extent of pyr-itization is related to concentration gradients of methane. In this model,

fractures and faults act as fluid conduits. We now assess such diageneticscenarios at Site C0002.

Magnetic parameter variations in Unit B (105–830.4 m CSF) ofIODP Hole C0002B (Fig. 3) reflect changes in mineralogy, especially thecomposition of iron sulphide minerals. Magnetic and SEM results in-dicate that greigite is the dominant magnetic mineral in the upper partof Hole C0002B. We have not identified pyrrhotite in the studied se-diments. Authigenic greigite with high susceptibility and saturationmagnetization can lead to magnetic enhancement of sediment (Robertset al., 1999; Dewangan et al., 2013), which contrasts with low sus-ceptibilities associated with paramagnetic pyrite and clays. Thus, therelative concentration of pyrite can be interpreted to increase with re-spect to greigite in the lower part of Unit B. Magnetic parameter var-iations do not correspond to lithological variations in Unit B (Fig. 3);thus, we relate the presence of greigite in these sediments to diageneticconditions associated with the presence of gas hydrates and AOM.

As indicated above, we have no samples from the gas-hydrate-bearing interval, so we use shipboard κ (Expedition 315 Scientists,2009; Strasser et al., 2014) as a proxy for magnetic mineral con-centration in this part of Unit B. Considering the approximately con-stant mineralogical bulk composition and density of Unit B (Expedition315 Scientists, 2009; Strasser et al., 2014), synchronous χ and κ fluc-tuations in the lower part of Unit B can be used to infer the magneticmineralogy in the gas-hydrate-bearing interval. The greigite-bearingsediments have relatively high susceptibility and SIRM/χ values, whichcorrelate well with each other (Figs. 3 and 7a), as indicated in a cross-plot of these parameters (Fig. 7b). Given this relationship, high volumemagnetic susceptibilities can be used as a proxy for greigite con-centration in Hole C0002K and C0002L. Thus, greigite is interpreted tooccur at depths of 200–390 m CSF based on sediment κ, which coincidewith depths at which gas hydrates are present. In situ gas hydrate sa-turation within the gas hydrate stability zone (GHSZ) has been esti-mated from different down-hole logging methods (Jia et al., 2017;Fig. 7a). The GHSZ corresponds to the region with high κ. A marked κdecrease occurs at the base of the GHSZ at ~390 m CSF in HoleC0002L. High κ values below the base of the GHSZ at ~410 to ~470 mCSF raise the question of whether they are indicative of the position of afossil GHSZ base and the former presence of gas hydrates, as proposedin other gas-hydrate-bearing sediments (e.g., Housen and Musgrave,1996; Larrasoaña et al., 2007). A paleo-BSR is observed in a bathy-metric notch near the seaward edge of the basin (Fig. 1b), where me-thane is interpreted to have been released (Bangs et al., 2010; Guoet al., 2013). The pink bars in Fig. 7a in the 100–640 m CSF interval areinterpreted to represent the present-day and fossil gas hydrate horizons.The presence of fossil GHSZs below 400 m CSF is consistent with theinterpreted dissociation of gas hydrate indicated by negative Cl− andpositive δ18O and δD of pore fluids at these depths (Toki et al., 2017).

Gas hydrate is sensitive to changes in pressure and temperature, andGHSZ migrations are likely to occur due to tectonic forcing (Fig. 7c).For example, we propose that the GHSZ was stable at ~1.2 Ma, andthat it then migrated in the sediment column due to faulting and/ormassive methane release and formed a new GHSZ (Fig. 7c). In Fig. 7a,we indicate with green bars zones in Hole C0002B with high con-centrations of normal faults and/or stratigraphic discontinuities. Theinterpreted base of the lowermost fossil GHSZ interpreted from mag-netic data occurs at ~640 m CSF (pink), which lies just above the up-permost faulted interval (green) indicated in Fig. 7a. This suggests thatfaults and fractures acted as pathways for methane migration as ob-served in the classic Hydrate Ridge sequence by Larrasoaña et al.(2007). This interpretation is consistent with evidence for sulphidic andmethanic gas and fluids at greater depth at this locality in the KumanoBasin (Ashi et al., 2007; Strasser et al., 2014).

5.4. Slope sediments

Magnetite is the dominant magnetic mineral in Unit C

M. Shi et al. Marine Geology 391 (2017) 1–12

9

(830.4–921.7 m CSF). This unit was deposited slowly above the car-bonate compensation depth and contains reworked material frommultiple sources in which organic matter was partially exhaustedduring initial deposition and early diagenesis (Underwood and Moore,2012; Ramirez et al., 2015). These factors limited iron sulphide for-mation during burial, which enabled magnetite preservation along withsignificant high coercivity mineral concentrations (probably hematite).

5.5. Relict sulphidic accretionary prism sediments

For sediments from Unit D (921.7–1052.5 m CSF), hysteresis loopsbefore slope correction are characteristic of the dominance of para-magnetic minerals (Fig. 4). Low SIRM and SIRM/χ values (Fig. 3)support this conclusion. Pyrite and rare PSD titanomagnetite are pre-sent. The dominance of paramagnetic minerals and the paucity ofdetrital or authigenic magnetic minerals suggest that pyritization re-actions proceeded to completion in this unit so that most detritalmagnetic minerals have undergone complete dissolution and any pre-cursor greigite has been transformed to pyrite. Methane sampled fromthe same location during IODP Expedition 338 has a mixed origin withthermogenic and oxidized microbial methane and enhanced sulphateconcentrations (Fig. 3), which suggest the existence of sulphidic andmethanic gas and fluids at greater depth (Ashi et al., 2007; Strasseret al., 2014). Pyrite formation along fractures (Fig. 5n, o) suggests thatthe fractures acted as conduits for methane migration, which has beenobserved elsewhere to cause later stage diagenetic magnetic mineralalteration (e.g., Larrasoaña et al., 2007). Thus, it appears that the ac-cretionary prism sediments of Unit D have undergone extensive sul-phidization, which is consistent with an enrichment of total sulphurirrespective of the approximately constant organic carbon content

(Fig. 3). In addition, the large numbers of faults and fractures in Unit D(Fig. 7a) both contributed to this sulphidization via late AOM andsupplied methane to overlying sediments, thereby contributing to gashydrate formation in overlying sediments.

6. Conclusions

Magnetic analysis of sediments from IODP Site C0002 has enabledconstruction of a detailed magnetic parameter sequence for the KumanoBasin since the late Miocene. Samples contain mixed magnetic mineralassemblages (magnetite, Ti-magnetite, greigite, and hematite) withdifferent dominant components in the different sedimentary and tec-tonic environments due to the varying effects of diagenesis, climatic,and tectonic forcing of sedimentation, and structural modification dueto local synsedimentary deformation. Ti-magnetite is the main mag-netic mineral in the studied sediments, except in the lower forearc basin(Unit B), where authigenic greigite dominates. In the forearc basin se-diments of Unit A, high initial detrital magnetic iron oxide concentra-tions in surface sediments were removed progressively with depth dueto diagenetic dissolution. Large titanomagnetite particles survived di-agenetic dissolution along with magnetic nanoinclusions within hostsilicates that protected the magnetic particles from sulphidic porefluids, and preserve a relict magnetic signal that records climaticmodulation of sediment delivery to Kumano Basin. Forearc basin se-diments of Unit B are dominated by magnetic signatures due to authi-genic iron sulphides (ferrimagnetic greigite and paramagnetic pyrite)that formed in association with gas hydrates and AOM. Heterogeneousgreigite concentrations are interpreted to be due to diagenetic varia-tions related to the presence of methane gas and hydrate. Methaneventing from deeper within the slope and accretionary prism sediments

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Fig. 7. (a) Down-core variations of magnetic parameters, including SIRM/χ, χ, and κ (Expedition 315 Scientists, 2009; Strasser et al., 2014), and gas hydrate saturation estimated fromlogging data using different methods (Jia et al., 2017). (b) Cross-plot of SIRM/χ and χ. (c) Summary tectonic cross-sections that illustrate the response of magnetic properties of thestudied sediments to migrations of the gas hydrate stability zone (modified from Moore et al., 2015).

M. Shi et al. Marine Geology 391 (2017) 1–12

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of Units C and D promoted formation of pyrite aggregates along sedi-ment fractures that contributed to further diagenetic depletion ofmagnetic signals in these units. The magnetic properties of the studiedsediments are consistent with the structural evolution of Kumano Basin,where deformation caused GHSZ migration through time so that thesediments of Site C0002 record both the present and fossil GHSZ po-sitions. The magnetic properties of Kumano Basin sediments provideuseful insights into the interplay between diagenetic, climatic, andtectonic processes that controlled the tectono-sedimentary evolution ofthis forearc basin.

Acknowledgements

Samples used in this study were provided by the Integrated OceanDrilling Program (IODP). We thank Dr. Lallan Gupta, staff, and tech-nicians at the Kochi Core Center, Japan, for their help. We also thank DrMichele Rebesco (Editor), Dr Fabio Florindo and an anonymous re-viewer for their careful review and constructive suggestions that sig-nificantly improved the paper. This work was supported by the NationalNatural Science Foundation of China (grants 41576062, 41422202,41688103 and 91128102), Fundamental Research Funds for the CentralUniversities (2652015297 and 2652015162), Gas Hydrate DatabaseProject (DD20160227), and Australian Research Council grantDP160100805.

Appendix A. Supplementary data

Supplementary data to this article can be found online at http://dx.doi.org/10.1016/j.margeo.2017.07.006.

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