Article
Volume 12, Number 8
23 August 2011
Q0AD16, doi:10.1029/2011GC003593
ISSN: 1525‐2027
Heat flow along the NanTroSEIZE transect: Results
from IODP Expeditions 315 and 316 offshore the Kii
Peninsula, Japan
Robert N. Harris
College of Oceanic and Atmospheric Sciences, Oregon State University, 104 COAS Administration
Building, Corvallis, Oregon 97331‐5503, USA ([email protected])
Friederike Schmidt‐Schierhorn
FB Geowissenschaften, Universität Bremen, Postfach 330 440, D‐28334 Bremen, Germany
Glenn Spinelli
Department of Earth and Environmental Science New Mexico Tech, Socorro, New Mexico 87801,
USA
[1] We report detailed thermal measurements undertaken during Integrated Ocean Drilling Program (IODP)
Expeditions 315 and 316 along the Nankai Trough Seismogenic Zone Experiment (NanTroSEIZE) transect
offshore the Kii Peninsula, Japan. Geothermal objectives included determining thermophysical rock properties of the cored material and characterizing the background thermal regime along this transect. New thermal conductivity measurements made with a divided bar are reported and supplement shipboard full‐ and
half‐space thermal conductivity measurements for a total of 938 thermal conductivity measurements. Thermal conductivity varies between about 1.0 W m−1 K−1 near the seafloor to 1.6 W m−1 K−1 at a depth of
1 km. Thermal conductivity generally increases with depth and correlates with variations in porosity
and lithology. Thermal gradients along the transect are characterized by 48 sediment temperature measurements from 6 sites. Thermal corrections for the effects of bathymetric relief and sedimentation improve the
confidence with which the heat flow values can be interpreted. Heat flow generally decreases with landward distance from the deformation front and varies from 70 mW m−2 just landward of the deformation
front to 54 mW m−2 at sites characterizing the outer fore‐arc high and to 57 mW m−2 at the Kumano Basin
Site. IODP heat flow measurements are significantly lower than nearby seafloor heat flow measurements.
This difference is most likely due to variations in bottom water temperature that have a large effect on
values of seafloor heat flow. Thus the heat flow of the Nankai accretionary prism is lower than previously
thought. We present thermal models of subduction along this transect and explore the impact of the initial
geotherm. Conductive plate cooling based on the age of subducting seafloor (20 Myr) under predicts the
observed heat flow. We find a good fit to the data using a geotherm appropriate for 10 Myr seafloor. The
extra heat is interpreted in terms of back‐arc thermal environments.
Components: 9700 words, 9 figures, 3 tables.
Keywords: NanTroSEIZE; heat flow; subduction.
Index Terms: 3035 Marine Geology and Geophysics: Midocean ridge processes; 3036 Marine Geology and Geophysics:
Ocean drilling; 3060 Marine Geology and Geophysics: Subduction zone processes (1031, 3613, 8170, 8413).
Received 2 March 2011; Revised 23 June 2011; Accepted 24 June 2011; Published 23 August 2011.
Copyright 2011 by the American Geophysical Union
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Harris, R. N., F. Schmidt‐Schierhorn, and G. Spinelli (2011), Heat flow along the NanTroSEIZE transect: Results from IODP
Expeditions 315 and 316 offshore the Kii Peninsula, Japan, Geochem. Geophys. Geosyst., 12, Q0AD16,
doi:10.1029/2011GC003593.
Theme: Mechanics, Deformation, and Hydrologic Processes at Subduction Complexes,
With Emphasis on the Nankai Trough Seismogenic Zone Experiment (NanTroSEIZE)
Drilling Transect
Guest Editors: D. Saffer, P. Henry, H. Tobin, and F. Chester
1. Introduction
[2] Knowledge of the temperature distribution of
accretionary prisms is important for understanding
physical and chemical processes such as fluid
flow, diagenesis, and the mechanics of faulting [e.g.,
Langseth et al., 1990; Yamano et al., 1992; Ruppel
and Kinoshita, 2000; Booth‐Rea et al., 2008].
Heat flow measurements provide the most direct
measure of the thermal state of accretionary prisms.
These measurements are most commonly based on
either seafloor probe determinations that measure
the in situ thermal gradient and thermal conductivity to depths of typically 4–5 mbsf, or borehole
temperature gradients that extend much deeper into
the sediment column, and thermal conductivities
measured on returned core material. Because of
their limited measurement depth seafloor probe
measurements are sensitive to shallow processes.
This effect is particularly manifest in dynamic
environments such as the overriding margins of
subduction zones where processes may significantly
change the thermal regime over short temporal and
spatial scales. These surficial variations are often
diagnostic of near seafloor processes such as variations in bottom water temperature, shallow deformation, and fluid flow but often obfuscate the deeper
thermal field. In contrast ocean drilling provides
access to deeper environments where downhole
tools, acoustic measurements, and logging technologies can provide important scientific insights into
the deeper thermal regime. The best chance of
obtaining a regional image of thermal processes
throughout a subduction margin is through a combination of numerous shallow seafloor probe measurements that capture spatial variability and deeper
but necessarily fewer borehole measurements that
are more representative of deeper processes.
[3] In this paper we report detailed thermal measurements undertaken during the Integrated Ocean
Drilling Program (IODP) Expeditions 315 and 316
to understand the thermal regime of the Nankai
accretionary prism offshore the Kii Peninsula.
These boreholes provided access for temperature
measurements to depths >100 mbsf and targeted
the frontal thrust area, the outer fore‐arc high where
a large splay fault intersects the seafloor, and the
Kumano Basin (Figure 1). Geothermal objectives
included determining thermophysical rock properties of the cored material and characterizing the
heat flow in this area. New thermal conductivity
measurements made with a divided bar are reported
and supplement shipboard full‐ and half‐space
measurements. Forty‐nine sediment temperatures
from 6 sites and close to 1000 thermal conductivity
determinations were used to calculate heat flow
along the NanTroSEIZE transect. These heat flow
values are then adjusted for the thermal effects of
bathymetric relief and sedimentation. We conclude
with a simple thermal model of subduction along
the transect.
2. Geologic Setting and Stratigraphy
[4] The Nankai Trough and accretionary prism are
the result of the Philippine Sea plate subducting
beneath Japan. The accretionary prism forms the
landward slope of the Nankai Trough and has
evolved since the Miocene. The prism is as much
as 100 km wide and consists of accreted and
underplated trench fill and Shikoku Basin sediments. The accretionary prism has been divided
into six structural zones: the trench, frontal thrust
fault zone, imbricate thrust zone, megasplay fault
zone, Kumano Basin edge fault zone, and the fore‐
arc basin [Moore et al., 2009].
[5] During IODP Expeditions 315 and 316, temperature and thermal conductivity measurements
were made in the region of the frontal thrust zone
(Sites C0007 and C0006), the megasplay fault zone
(Sites C0008, C0004, and C0001) and in the
Kumano Fore‐arc Basin (Site C0002). The topographic setting and stratigraphy for each region is
briefly described to place the thermal data in context. Because clay and quartz content influence
thermal conductivity we emphasize this aspect of
the stratigraphy.
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Figure 1. The NanTroSEIZE drilling transect and heat flow. (a) Seismic reflection line with drill site locations.
Figure modified from Moore et al. [2009]. Inset shows regional tectonic map and location of the seismic line
(red). FSC is fossil spreading center and PSP is the Philippine Sea plate. (b) Heat flow around the NanTroSEIZE
transect. Diamonds show IODP Sites, circles show seafloor heat flow sites, triangle show older data, and squares show
long‐term temperature monitoring sites (H. Hamamoto et al., submitted manuscript, 2010). Symbols are color coded
by heat flow value. All heat flow values have been corrected for the thermal effects of bathymetry as described in the
text. IODP heat flow values have also been corrected for the effects of sedimentation.
[6] Sites C0007 and C0006 are within the frontal
thrust zone on a topographic high that forms the
northeast headwall of a landslide. C0007 is on the
slope up to the high and C0006 is on the slope
break (Figure 1). Sediments recovered at these sites
reflect underthrust and accreted material overlain
by slope deposits. The stratigraphy of Sites C0007
and C0006 is described by Kinoshita et al. [2009]
and Screaton et al. [2009]. Four units were cored at
Site C0007; the upper three of these were also
cored at Site C0006 (Figure 2). Unit I comprises
the slope deposits and is interpreted as being
deposited by hemipelagic settling, turbidity currents,
and soft sediment slumping on an oversteepened
slope. It is composed of nannofossil‐bearing mud,
interbedded sand layers, and minor volcanic ash.
Unit II is interpreted as accreted trench wedge
material; it consists of fine‐grained mud to sand‐
and gravel‐rich deposits that coarsen upward. At
the base of unit II a significant age transition
from Pleistocene and younger sediments above to
Miocene (2.5 Ma and older) below marks the
boundary with the Shikoku Basin facies (unit III).
Unit III consists of mud with ash and tuff layers,
interpreted to have been deposited by hemipelagic
settling in a deep marine basin. At Site C0007 the
base of unit III was cored and is marked by evidence of a fault zone [Kinoshita et al., 2009]. Unit
IV consists of sands and recovery was poor. At
both sites, clay content increases with depth from
∼40% to 60% and quartz content decreases somewhat with depth.
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Figure 2. Lithology and thermal conductivity data.
[7] Sites C0008, C0004, and C0001 are in the
megasplay fault zone, just southeast of a prominent
ridge that forms the seaward boundary of the
Kumano Basin (Figure 1). The stratigraphy of
these sites is described by Kinoshita et al. [2009]
and Kimura et al. [2011]. Sediments recovered
from Site C0008 reflect upper slope basin deposits
(Figure 2). Two units were identified. Unit I consists of silty clay with a substantial component of
calcareous nannofossils, siliceous biogenic debris,
and volcanic ash. Lower in the unit, interbedded
mud clast gravels and silty clay beds dominate.
Unit II consists of sand‐rich turbidites with pebbly
sandstone and minor silty clay interbeds. Clay
mineral content varies between ∼25 and 55%;
quartz content is relatively uniform at ∼20%. Site
C0004 targeted the conspicuous splay fault at
300 mbsf and recovered accretionary prism sediments in the hanging wall of the splay fault, slope
sediments that overlie the prism, and slope basin
sediments in the footwall of the splay fault. Four
units were identified. Unit I consists of hemipelagic
nannofossil‐rich mud with rare volcanic ashes and
rare thin sands and is interpreted as slope sediment.
In the uppermost portion of unit II sedimentary
breccias and hemipelagic muds dominate while in
the lower portion hemipelagic muds with rare turbidite sands dominate. These sediments represent
mass wasting deposits and prism sediment, respectively. Unit III consists of ash‐bearing hemipelagic
muds with scattered volcanic ashes and unit IV
consists of hemipelagic mud with turbidite sands
and rare volcanic ashes. In general, clay content
increases with depth from 40 to 65% through unit
III and then decreases to about 45% in unit IV.
Quartz content is relatively uniform at about 20%.
Site C0001 documented the hanging wall of the
megasplay and intersected two units. The upper
section of unit I consists of nannofossil‐bearing
hemipelagic mud with volcanic ash, and the lower
portion contains sand and himipelagic mud and is
interpreted as reflecting slope apron facies. Unit II
is dominated by hemipelagic mud representing the
upper accretionary prism. Clay minerals increase
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dramatically through unit I from about 30 to 60%
and are relatively constant through unit II. Quartz
content is relatively uniform at approximately 20%.
[8] Site C0002 is in Kumano Basin and targeted
the sedimentary sequence of the upper basin
(Figure 2). Four units were cored and consist of
hemipelagic muds with fine sands and silt turbidites in the upper sections. Clay content shows
small variations with depth and varies between
about 30 to 65 with a significant decrease in unit
IIIa. Quartz generally decreases with depth and
varies between about 15 and 35%.
3. Thermal Conductivity
[9] To understand the thermal structure of the
Nankai prism, it is important to have a good
knowledge of the thermophysical rock properties.
Of these properties thermal conductivity is of primary importance because it exerts a fundamental
control on the distribution of temperatures. Thermal conductivity varies as a function of porosity
and mineralogy and can vary by a factor of three or
more from 0.9 W m−1 K−1 for porous clays to
4.5 W m−1 K−1 for lithified quartz‐rich sandstones
[e.g., Zoth and Haenel, 1988].
[10] We measured thermal conductivity using three
different techniques. In the upper portions of
Holes, sediments were generally unlithified and
full‐space measurements were made [Von Herzen
and Maxwell, 1959]. In the deeper portions of
Holes, where sediments were lithified and pieces
were big enough (>5 cm), we used the half‐space
technique [Vacquier, 1985]. Both of these methods
use a needle probe that contains a thermistor and
heating wire that approximates a line source. In the
full‐space method the needle is inserted through a
small hole in the core liner into the unlithified
sediment, and in the half‐space method the needle
is insulated within an epoxy disk but exposed on
one face. The face of the disk with the exposed
needle is placed against the face of lithified material being measured. The needle is aligned parallel
to the core axis. Supplementary measurements
made perpendicular to this direction show no evidence of anisotropy in the lithified material. Both
the full‐ and half‐space methods are transient
techniques that approximate the heating element as
an infinite line source. The needle is heated with
constant power and the temperature time response
yields the thermal conductivity. The experimental
uncertainty is estimated to be 5%. Many of the full‐
10.1029/2011GC003593
and half‐space measurements were made in triplicate, and repeated measurements were usually within
1 to 2% of the mean. Both methods were calibrated
with a rubber standard (0.517 W m−1 K−1), and two
glass standards (1.237 and 1.623 W m−1 K−1). These
calibrated values showed a small offset relative to
the standard. Secondary calibration constants were
determined based on a least squares inversion
between these measured values and standards
[Kinoshita et al., 2009]. These adjustments increased
the measured values by approximately 4%.
[11] In many of the Holes, sections of returned core
material were too lithified to insert a needle but so
friable that pieces were too small to measure with
the half‐space disk. In these regions lithic fragments were collected and thermal conductivity was
measured using a divided bar apparatus [e.g., Sass
et al., 1971, 1984]. The divided bar is a steady
state technique in which a heat flux is applied across
a sample and the thermal gradient is measured.
Samples are placed in a cylindrical cell and vacuum
saturated with water to remove air. The bulk thermal
conductivity, kb, of the cell is determined as the ratio
of the applied heat flux. The grain thermal conductivity is determined from the geometric mean
kb ¼ kw kg1
ð1Þ
where kw is the thermal conductivity of water (0.6 W
m−1 K−1), kg is the grain conductivity, and f is the
porosity of the cell. The bulk formation conductivity
is then determined using equation (1) a second time
by using the just determined grain conductivity and
formation porosity. All divided bar measurements
were made on a modified tandem apparatus at the
University of Utah [Chapman et al., 1981] that was
calibrated with both fused quartz and crystalline
quartz. Experimental uncertainties for these measurements are estimated to be 3%.
[12] Eleven half‐space and divided bar measure-
ments were made on samples within 10 cm of each
other. These measurements agree to 5%, although
the standard deviation is approximately 13%.
Comparisons between five full space and divided
bar measurements have a mean agreement and
standard deviation of 2 and 22%. The comparisons
show that the methods generally give similar
results with no systematic offsets.
[13] All thermal conductivity measurements were
corrected for in situ pressure and temperature.
Thermal conductivity increases with pressure; for
each 1800 m of depth the thermal conductivity is
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either too lithified to get good contact with the
needle for full‐space measurements, or pieces too
small to avoid edge effects with the half‐space
measurements. The generally greater scatter at very
shallow depths may be due to low values attributed
to gas expansion and/or core disturbance.
[15] In general thermal conductivity varies between
Figure 3. Thermal conductivity as a function of porosity and color coded by drill site. Note the general inverse
correlation between thermal conductivity and porosity as
well as a geographic distribution that primarily reflects
quartz content.
about 1.0 W m−1 K−1 near the seafloor to 1.6 W
m−1 K−1 at depth. Measurements may be biased
toward low values because of core disturbance and
greater recovery of muds instead of sands relative to
their in situ occurrence. Bulk thermal conductivity
is inversely proportional to porosity (Figure 3).
This correlation is most pronounced at sites C0007,
C0004, C0001 and C0002 where variations in
porosity are significant. Additionally, second‐order
variability is present that can be understood in
terms of lithology. High values tend to be associated with sands and higher quartz content while
low values tend to be associated with clays and ash
layers. For example, the frontal thrust sites (C0007
and C0006) have a higher fraction of quartz content
and are associated with generally high thermal
conductivities. In other sections where muds or
volcaniclastics dominate, thermal conductivity
values are low (C0001 and C0002).
4. Temperature Measurements
[16] Equilibrium temperatures were measured using
increased by 1% [Ratcliffe, 1960]. In porous rocks
temperature influences thermal conductivity in two
competing ways. The thermal conductivity of rock
matrix is inversely related to temperature [Zoth and
Haenel, 1988], whereas the thermal conductivity of
water increases with temperature [Keenan et al.,
1978]. We apply a temperature correction of +1%
for each +20°C change in temperature between the
laboratory and estimated in situ temperatures, a
value intermediate between +5% suggested by
Ratcliffe [1960] for a high‐porosity, water‐saturated
sediment and mean value of −3% derived from data
reported by Clark [1966] for several hard rocks.
Thermal gradients determined at each site are discussed in section 4. These corrections increased
values of thermal conductivity by 1–2%.
the APCT‐3 and SET temperature tools. The
APCT‐3 tool fits within the cutting shoe of the
hydraulic piston core used for coring sediments.
The inner core barrel is shot 9.5 m into sediments at
the bottom of the hole, sufficiently deep to avoid
the thermal influence of drilling. After the core
barrel penetrates, it is decoupled from the drill
string and left in place from 7 to 10 min. During
this period a temperature‐time series of the
decaying heat pulse from the frictional of penetration is recorded. These temperatures are extrapolated to the equilibrium temperature using a tool
response function [Heesemann et al., 2007]. The
SET tool measures temperatures in stiffer sediments and requires a wireline trip. This tool
extends 1.4 m below the drill bit and is held in
place with weight applied by the driller.
[14] Thermal conductivity measurements are shown
[17] Forty‐eight high‐quality in situ temperature
in Figure 2. In some Holes, scatter is greatest near
the deepest full‐space measurements (e.g., C0007
and C0004) or the shallowest half‐space measurements (e.g., C0002). This scatter likely reflects
applying the techniques to questionable material,
measurements were made at seven different sites
on Expeditions 315 and 316 (Figure 4a). Each
temperature‐depth profile is based on at least five
equilibrium temperatures (Holes C0007 and C0004)
and up to 10 equilibrium temperatures (Hole
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Figure 4. (a) Equilibrium temperatures plotted as a function of depth (open symbols) and thermal resistance (solid
symbols). Temperature is plotted on a relative scale to avoid overlap. Heat flow (q) and bottom water temperature
(BWT) are given. Solid lines show best fit to data as a function of thermal resistance. (b) Residual temperature is
the difference between the equilibrium temperature and the best fitting linear regression.
C0008C). The deepest equilibrium temperature is
approximately 250 m below seafloor (mbsf) at
Hole C0008A, but most Holes have temperatures
extending to only 150 mbsf because of increasing
sediment induration. Equilibrium temperatures are
plotted as a function of depth and thermal resistance
on a relative scale to avoid overlap (Figure 4a).
Subbottom temperatures can be expressed as
[Bullard, 1939]
T ð zÞ ¼ To þ qo
N
X
Dzi
k
ð zÞi
i¼1
ð2Þ
where To is the bottom water temperature, qo is the
background heat flux, ∑Dzi /k(z)i is thermal resis7 of 18
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Table 1.
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Best Linear Fit Parameters
Site
Number of
Measurements
BWT
(°C)
Heat Flowa
(mW m−2)
Standard Deviation
of Fit (m2 W−1 K−1)
Heat Flowb
(mW m−2)
Heat Flowc
(mW m−2)
C0007
C0006
C0008A
C0008C
C0004
C0001
C0002
5
6
9
10
5
7
6
1.6
1.6
2.0
2.2
1.8
2.0
2.4
60
36
51
55
54
46
41
0.1
0.1
0.5
0.2
0.1
0.3
0.1
63
43
48
50
51
53
44
70
45
51
53
54
60
57
a
Heat flow computed from Bullard analysis.
Heat flow corrected for the effects of bathymetry.
Heat flow corrected for sedimentation.
b
c
tance, Dzi is the depth spacing and k(z)i is the thermal conductivity corresponding to that depth spacing. In general thermal resistances appear linear and
are in general well fit by a constant gradient suggesting constant heat flow. The standard deviation
of the fit to the thermal resistance is less than or equal
to 0.5 m2 W−1 K−1 (Table 1). Residuals are plotted in
Figure 4b and give a better sense of discrepancies
between individual thermal resistances and best fitting lines. Holes C0008A and C0008C are notable
because they are only separated by ∼200 m. Temperatures in Hole C0008A are generally more noisy
than those in Hole C0008C but the heat flow differs
by only 5 mW m−2, or about 10%, indicating internal
consistency. The reason for the increased noise at
Hole C0008A is unknown.
[18] Values of heat flow range from 60 to 36 mW m−2
(Table 1) and in general decrease with distance
landward of the deformation front, consistent with
subduction. The heat flow value at Hole C0006
appears anomalously low.
5. Corrections to Heat Flow
Determinations
5.1. Bathymetric Correction
[19] Bathymetry has the potential to distort the
thermal field by increasing the thermal gradient
under bathymetric lows and decreasing it under
bathymetric highs. High‐resolution multibeam
bathymetry with a grid size of approximately 100 m
[Ike et al., 2008] shows pronounced relief at the
drilling sites (Figure 1). Our strategy for correcting
the thermal field is based on techniques formulated
by Blackwell et al. [1980]. For each borehole we
construct a grid of seafloor temperatures based on
three‐dimensional bathymetry and CTD casts from
the world databank. Seafloor temperatures are
projected onto a horizontal plane defined at the top
of the borehole using the best fitting thermal gradient. Finally this surface is fit with a Fourier series
and upward (depth positive down) continued to
generate the thermal perturbations created by the
bathymetry. Because the best fitting gradient used
in the projection step is also the unknown being
solved for, this scheme is necessarily iterative. The
bathymetrically induced gradients are subtracted
from the observed gradient to generate the background thermal gradient.
[20] Heat flow values have been corrected for
bathymetry and are reported in Table 1. Heat flow
increases for Sites C0007, C0006, C0001, and
C0002 consistent with being sited on local ridges
and topographic highs (Figure 1). Heat flow
slightly decreases for Sites C0008A and C0008C,
and C0004 consistent with being sited in local
topographic lows. The largest change is associated
with Sites C0006 (+19%) and C0001 (+16%) that sit
on a slope break and local seafloor high, respectively. These large changes show the importance of
this correction in regions of large relief.
5.2. Sedimentation Correction
[21] Sedimentation and subsidence can transiently
depress surface heat until the deposited sediments
are warmed to the background equilibrium temperature. Similarly, unroofing of sediments through
erosion or mass wasting events can transiently
increase the surface heat flow until the unroofed
section of sediments is cooled to the background
equilibrium temperature. Thrust faults, normal
faults, biostratigraphic and paleostratigraphic age
reversals, and porosity data indicate a combination
of these mechanisms have likely perturbed the
thermal field. The linearity of the temperature‐
depth profiles suggests that these processes do not
dominate the current thermal field, although the
vertical spacing of temperature measurements
degrades resolution to these events. Here we focus
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Figure 5. Deposition data for (a) frontal thrust area, (b) megasplay fault zone, and (c) Kumano Basin. Symbols are
based on biostratigraphy and paleomagnetic data, and solid lines show sediment accumulation rates used in sedimentation correction. Inset shows porosity data and best fitting curves.
on the most recent large events because these will
have the biggest impact on the heat flow. Our
sedimentation correction is based on the work by
Hutchison [1985] as implemented in a one‐
dimensional finite difference algorithm [Wang and
Davis, 1992] that includes the thermal effects of
variable sedimentation rates, compaction, and different thermal properties between the sediment and
basement rocks. This correction is most sensitive to
the combination of sediment thickness and its rate
of deposition [Hutchison, 1985; Wang and Davis,
1992; Martin et al., 2004].
[22] Sedimentation rates and intervals of erosion at
these sites, particularly in the megasplay fault zone is
episodic due to irregular periods of fault and structural activity, as well as slumping and mass transport
deposition [Strasser et al., 2009, 2011]. Sediment
accumulation rates were calculated from biostratigraphic and paleomagnetic data documenting sedi9 of 18
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Table 2.
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Sedimentation Correction
Site
Initial
Porosity
Compaction
Constant (km)
Durationa
(Myr)
Accumulation
Rate (m/Myr)
Duration
(Myr)
Accumulation
Rate (m/Myr)
Duration
(Myr)
Accumulation
Rate (m/Myr)
C0007
C0006
C0008A
C0008C
C0004
C0001
C0002
0.50
0.50
0.61
0.61
0.61
0.63
0.62
1.06
1.85
1.09
1.09
1.09
1.99
2.07
0.5
0.6
1.8
1.8
2.5
2.0
1.0
24
40
27
27
30
100
140
1.0
1.2
0.5
0.5
2.0
2.0
1.0
480
150
520
520
140
13
1300
–
–
1.7
1.7
–
2.0
5.0
–
–
35
35
–
320
140
a
Duration of sedimentation beginning with the most recent phase.
ment accumulation [Kinoshita et al., 2009]. Representative accumulation curves for each of the drilling
areas are shown in Figure 5. We have fit the accumulation rate data with two or three linear trends and
the durations and rates are given in Table 2.
[23] We convert accumulation rates, Vo, to sedi-
mentation rates, Vs, through [Hutchison, 1985]
Vs ¼
V o ð1 o Þ
ð1 ð zÞÞ
ð3Þ
where fo is the initial porosity and f is the porosity
as a function of depth, z. Porosity curves were fit to
porosity data at each borehole using Athy’s law
z
ð zÞ ¼ o exp L
ð4Þ
where L is an empirically derived constant related
to compaction. Porosity data for each region
(Figure 5) are best fit in a least squares sense and
best fitting parameters are reported in Table 2. In
general the best fitting porosity trends fall in a
small range with the exception of sites in the frontal
thrust region. The initial porosity associated with
C0007 and C0006 is 50% and suggest that this
material undergone compaction by material that
has been removed [Screaton et al., 2009]. At the
other sites porosity does not decrease to <50% until
150 or 200 mbsf in the slope region or Kumano
Basin, respectively. The erosion or removal of
material would increase the surface heat flow, but
because heat flow values at these sites appear relatively low, the transient effects of erosion may
have diminished.
[24] Slope sites C0001, C0004, and C0008 have
been subject to recent erosion, with surface hiatuses
at 0.30, 0.44, and 0.06 Ma [Kinoshita et al., 2009].
These erosive events are likely due to slumping and
mass wasting [Strasser et al., 2011]. Corrections
for slumping to marine heat flow data are seldom
applied because heat flow observations mostly try
to avoid these areas. Here we assess the impact of
these surface hiatuses on heat flow by modeling the
thermal impact in terms of an instantaneous step
change in seafloor temperature, T(z, t), expressed as
z
T ð z; Þ ¼ DTerf pffiffiffiffiffiffiffiffi
4
ð5Þ
where z is depth, t is time before present that the
event occurred, DT is the difference between the
original seafloor temperature and the new seafloor
temperature due to unroofing, erf is the error
function, and a is thermal diffusivity. Equation (5)
shows that the impact of slumping events depends
the magnitude of unroofing, through DT and z, and
the time of the event in the past, t. We compute
percent changes in heat flow as a function of
unroofing thickness and time since the event
(Figure 6). For these perturbations we used a
background thermal gradient of 0.06°C/m, a thermal conductivity of 1 W m−1 K−1. Near subbottom
perturbations are much larger than indicated by the
figure; we fit the temperature profile in a least
squares sense over a depth range of 0 to 100 m to
assess its impact on our heat flow determinations.
For this set of parameters, largest perturbations
correspond to unroofing events between 10 and
100 years before present. After approximately
10,000 years the thermal effects have attenuated.
Extending the depth range of the calculation
extends the time period over which the gradient is
sensitive to erosive events.
[25] All three sites have relatively modest sediment
accumulation rates in the near past (Figure 5), and
we explore the thermal impact of mass wasting at
Site C0004 where the hiatus is the longest. A
sediment accumulation rate of 0.03 mm/yr yields a
thickness of approximately 13 m. This may add an
approximately 10% increase to the observed heat
flow provided the mass wasting effect occurred
between 10 and 100 years before present.
[26] Sedimentation corrected heat flow values are
given in Table 1. Sedimentation corrections on the
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Figure 6. Percent change to thermal gradient based on unroofing events as a function of time and thickness.
prism slope vary between approximately 5 and
10%. Site C0002 within the Kumano Basin has a
sedimentation correction of 22% consistent with
rapid Pleistocene sedimentation.
6. Discussion of Heat Flow
Observations
[27] Figure 7 shows the new corrected heat flow
results in the context of the uncorrected measurements and existing data along the NanTroSEIZE
transect. Including both the bathymetric and sedimentation correction leaves the heat flow at Site
C0006 anomalously low. The source of this low
value is not known. It may be that bottom water
infiltrated the lithology during the temperature
measurement, although temperature‐time series of
each measurement appear reasonable [Expedition
316 Scientists, 2009]. Neglecting the anomalously
low value at Site C0006, the borehole values of
heat flow generally decrease with distance from the
deformation front consistent with the downward
advection of heat by the subducting plate.
Figure 7. Heat flow data along NanTroSEIZE transect.
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[28] It is interesting to compare nearby seafloor
measurements of heat flow with the borehole
values of heat flow. We project existing heat flow
from shallow probe (4.5–6 m) measurements
(H. Hamamoto et al., Heat flow distribution and
thermal structure of the Nankai Subduction zone
off the Ki‐I Peninsula, submitted to Geochemistry
Geophysics Geosystems, 2010) onto the profile
from distances as great as 10 km. The majority of
these stations are within 2 km of the transect and lie
near sites C0004 and C0008 (Figure 1). We focus
on these data and for comparison we have also
corrected these data for the effects of bathymetry
using an assumed maximum measurement depth of
5 m. Reasonable variations in this parameter have a
negligible influence on the corrected heat flow
value. With the exception of two values at a distance of about 8 km from the deformation front,
this correction decreases observed heat flow values
on average by less than 5%. These other two values
sit on a narrow trough and have an exceptionally
large correction that decreases their magnitude by
about 40%. It is interesting to note that the bathymetric correction for borehole heat flow systematically increase values, whereas for shallow probe
data, values are systematically decreased. This may
be due to placing probe measurements in basins or
trough where the probe is more likely to penetrate
sediments, and placing boreholes in basement
highs or slopes. We have not corrected probe heat
flow values for the effects of sedimentation but
based on the correction at sites C0004 and C0008,
the effect would increase these values by about
10%. Heat flow values for the shallow probe data
are based on an average of in situ thermal conductivity determinations from the shallow probe
measurements of 1.04 W m−1 K−1. This value is
nearly identical to thermal conductivity measurements at C0004 and C0008 that are less than
10 mbsf, indicating that the difference in heat flow
is due to the difference in thermal gradients. The
shallow probe measurements have a mean and
standard deviation of 73 and 11 mW m−2, respectively, 15 mW m−2 greater than the borehole heat
flow measurements that have a mean and standard
deviation of 58 and 8 mW m−2, respectively,
excluding Site C0006.
[29] It is difficult to unequivocally attribute the
discrepancy between these two data sets to a particular cause or process because downhole measurements and shallow probe measurements have
uncertainties that are both unique and common.
Accretionary prisms are dynamic environments and
discrepancies may be related to real issues of spa-
10.1029/2011GC003593
tial heterogeneity. However, because these three
borehole measurements and nine shallow probe
measurements are systematically offset from each
other it seems reasonable to assume that this
reflects curvature in the temperature field of the
subbottom. Two processes not already considered
that could produce this curvature include the upward
advection of heat by fluids and bottom water
temperature variations.
[30] The upward flux of fluids could cause the
appropriate curvature. Low chloride cold seeps and
chemosynthetic communities are observed southeast of the transect where splay faults intersect the
seafloor [Toki et al., 2004]. Goto et al. [2008]
measured heat flow in a biological community and
found heat flow in excess of 180 mW m−2, much
higher than the seafloor heat flow measurements
near the transect. We solve the one‐dimensional
steady state advection‐diffusion equation subject to
a constant seafloor temperature and constant gradient at depth and express the mean vertical fluid
velocity, Vz, as a function of the surface heat flow,
qo, [e.g., Harris et al., 2010]
Vz ¼
kf
kf G
ln
w cw H
qo
ð6Þ
where f is the porosity, rw cw is the heat capacity of
water, H is a length scale, and G is the thermal
gradient at depth. An upward fluid flux of
approximately 50 mm/yr would be required to
increase the subbottom heat flow by 15 mW m−2.
However, pore water analysis made on returned
core material is consistent with in situ diagenetic
reactions and there is no clear evidence of fluid
flow along the decollement and associated splay
faults at the sites drilled during Expeditions 315
and 316 [Solomon et al., 2008; Wheat et al., 2008].
These results suggest that advective fluid flow is
unlikely responsible for the difference between
seafloor and deeper values of heat flow in this area.
[31] Bottom water temperature variations can per-
turb shallow thermal gradients and has been suggested as an explanation of the scatter in heat flow on
the accretionary prism slope (H. Hamamoto et al.,
submitted manuscript, 2010). We summarize their
argument and explore its implications for reconciling the seafloor probe measurements with borehole
measurements of heat flow. The Nankai Trough is
affected by the Kuroshio current, a strong western
boundary current that imparts bottom water temperature variations to the shallow subbottom
[Yamano et al., 1984; Hamamoto et al., 2008]. The
seafloor probe measurements were made at bathy12 of 18
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Table 3.
HARRIS ET AL.: HEAT FLOW ALONG THE NANTROSEIZE TRANSECT
[32] Together these results suggest that heat flow
Model Parameters
Unit
Incoming sediment
Oceanic crust
Accreted sediment
Continental crust
Mantle
Heat
Heat
Thermal
Capacity
Conductivity Production
(W m−1 K−1) (mW m−1) (MJ m−1 K−1)
1.2
2.9
1.5–2.5a
2.5
3.1
10.1029/2011GC003593
1.9
0.02
1.9–2.4a
0.4–2.4b
0.02
2.5
3.3
2.5
2.5
3.3
a
Thermal conductivity and heat production decrease linearly from
the prism toe to the backstop.
b
Heat production decreases linearly from 0 to 10 km depth.
metric depths that range between approximately
2700 to 3050 m. Although there are no long‐term
records of bottom water temperature variations at
these depths, bottom water temperatures at a depth
of 2525 m were recorded over a two and half year
period and shows annual peak‐to‐peak temperature
variations of about 0.3°C [Hamamoto et al., 2008].
This variation can produce a maximum temperature
change of 0.115°C at 5 mbsf; corresponding to an
apparent change in heat flow of about 23 mW m−2,
large enough to reconcile the probe measurements
with the borehole measurements of heat flow. The
majority of these measurements, six out of nine,
have a penetration of only 2 m (H. Hamamoto et
al., submitted manuscript, 2010) that can lead to
greater perturbation than that calculated above.
There is an additional clue that the discrepancy
between these sets of data may be due to bottom
water temperature variations. The shallow probe
data come from three cruises (Figure 7). We note
cruises KT‐05 and KT‐06 form two distinct clusters
suggesting that the heat flow values may in part a
function of cruise timing. The means of these two
data sets are 63 and 82 mW m−2, a difference of
19 mW m−2. The standard deviations of the two
subsets are 4 and 11 mW m−2, respectively, values
less than the difference between their means. Two
additional values are from a third cruise KR‐04
overlap both of the clusters. Together this evidence
suggests that part of the variation in heat flow is a
function of cruise date and may therefore be due to
variations in bottom water temperature. Cruise
KT‐05 was 25 February through 3 March 2005,
and cruise KT‐06 was later in the spring 11 April
through 22 April 2006 (M. Yamano, written communication, 2011). It is interesting to note that heat
flow at Site C0007 lies on a trend between the heat
flow values collected on cruises KR‐02 and KR‐04.
These data at greater water depths (>3 km) suggest
that significant bottom water temperature variations
may not extend to these depths.
on the accretionary prism along the NanTroSEIZE
transect may be lower than previously estimated for
two reasons. First bathymetric corrections generally lower seafloor probe values of heat flow. This
may reflect a bias toward placing these measurements in troughs where there is a greater likelihood
of probe penetration. Second, because the borehole
temperature measurements come from deeper in the
prism and are therefore much less susceptible to
bottom water temperature variations, we suggest
that heat flow determinations for the borehole sites
are more reflective of the thermal state of the
accretionary prism.
7. Thermal Models
[33] The thermal state of the shallow subduction is
primarily a function its convergence rate and dip,
the thermal properties of the margin, and importantly the temperature structure of the incoming
plate [e.g., Cloos, 1985; Dumitru, 1991]. We use a
transient 2‐D finite element model [Wang et al.,
1995] to simulate temperatures in a cross section
along the NanTroSEIZE transect. Heat is transferred advectively with the subducting plate and
conductively through the fore arc. The Philippine
Sea plate is subducting at a rate of 4 cm yr−1 [Seno
et al., 1993]. The geometry of the model is derived
from seismic reflection data for the shallow portion
of the subduction zone [Moore et al., 2009] and
tomographic imaging for the deeper slab geometry
[Hirose et al., 2008]. We parameterize the geometry of the model in terms of 5 units (Table 3).
Model parameters are based on a combination of
NanTroSEIZE drilling results [Kinoshita et al.,
2009] and parameters used along the Muroto transect [Hyndman et al., 1995; Wang et al., 1995;
Spinelli and Wang, 2008].
[34] The upper boundary for the model has a fixed
temperature of 0°C, and the base of the subducting
plate is assigned a temperature of 1400°C. The
thermal regime of the subduction thrust is insensitive to this basal boundary condition. The landward boundary is also placed sufficiently far from
the seismogenic zone to avoid boundary effects. At
the landward boundary, we prescribe a surface heat
flow of 70 mW m−2 and an adiabatic gradient
through the mantle wedge. The thermal effects of
mantle wedge flow are coupled to the subducting
slab using an isoviscous mantle rheology [Peacock
and Wang, 1999]. The details of the mantle rheology have only a small effect on the seismogenic
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Figure 8. Heat flow along the strike of the Nankai
Trough and seaward of the accretionary prism. The location of the fossil spreading center (FSC) and NanTroSEIZE transect are shown. Solid line shows conductive
predictions for half‐space cooling based on a half spreading rate of 30 mm yr−1 [Sdrolias et al., 2004]. The dashed
line is depressed 20% to account for sedimentation.
Tick marks on the line showing the position of the
NanTroSEIZE transect show the predicted half‐space
cooling values for 10 and 20 Ma oceanic crust. Figure
modified from Yamano et al. [2003].
portion of the subduction thrust [Currie et al.,
2002].
[35] The temperature structure of the incoming
Philippine Sea plate plays an important role in
the model outcome but also introduces significant
uncertainty. There are at least two factors that
increase this uncertainty here: (1) the tectonic history is transient and may impact the temperature
field [Wang et al., 1995], and (2) the Philippine Sea
plate was formed as a back arc to the Izu‐Bonin
subduction zone and the thermal structure of back
arcs are not well known in detail [Peacock, 2003;
Currie and Hyndman, 2006]. Both of these factors
are discussed below.
7.1. Tectonic History of Subduction
[36] The Philippine Sea plate formed as a back‐arc
basin to Pacific plate subduction at the Izu‐Bonin
Trench (Figure 1, inset). Back‐arc spreading originated along a now fossil spreading center located
offshore Cape Muroto and marked by the Kinan
seamount chain [Kobayashi and Nakada, 1978;
Chamot‐Rooke et al., 1987; Okino et al., 1994].
Sdrolias et al. [2004] analyzed the tectonic history
of the Shikoku Basin based on magnetic, gravity,
and bathymetric and the salient features are briefly
10.1029/2011GC003593
summarized here. The Shikoku Basin originated at
approximately 28 Ma due to a propagating rift in
response to slab rollback of the Izu‐Bonin Trench.
At this time the Izu‐Bonin Trench was south of
22°N and old (>75 Ma) Pacific plate was subducting at the position of the Kii Peninsula. Contemporaneously with slab roll back of the Pacific
plate, the Izu‐Bonin Trench migrated to the northeast such that the trench was approximately at the
position of the Kii Peninsula transect at 14 Ma.
Rifting of the Shikoku Basin occurred between 28
and 15 Ma. Since about 14 Ma, when the Izu‐Bonin
Trench migrated past the position of the Kii Peninsula,
the Nankai Trough has subducted back‐arc material.
Currently back‐arc seafloor with a magnetic age of
20 Ma is being subducted.
7.2. Regional Heat Flow and Thermal
Structure of Shikoku Basin
[37] Currie and Hyndman [2006] surveyed obser-
vations of back‐arc basins that bear on their thermal state and found that nearly all back‐arc basins
are anomalously warm even in the absence of
recent extension. They also found that the thermal
anomaly extended over wide areas of the back arc
(250 to >900 km). Thus we expect the thermal state
of the incoming Shikoku Basin may be anomalously warm.
[38] The Nankai Trough and Shikoku Basin have
been the focus of heat flow studies for decades [e.g.,
Watanabe et al., 1970]. Because oceanic heat flow
is most often understood in terms of conductive
cooling models as a function of age, Yamano et al.
[1992, 2003] viewed heat flow perpendicular to
the fossil spreading center. Heat flow is a maximum near the fossil spreading ridge offshore Cape
Muroto (Figure 8). Although heat flow generally
decreases with distance from this fossil spreading
center both westward toward Cape Ashizuri, and
eastward toward the Kii Peninsula, it is significantly higher than conductive cooling curves based
on seafloor age [Kinoshita and Yamano, 1995;
Kinoshita et al., 2008]. Given the scatter in the
data, the heat flow can also be approximated by a
constant value. Offshore the Kii Peninsula, and
seaward of the Nankai Trough, seafloor values of
heat flow are about 110 mW m−2. The cooling
effect of sedimentation is thought to transiently
depress this value by about 20%, putting the equilibrium heat flow at about 130 mW m−2. Usingphalf‐
ffiffiffiffiffiffiffiffi
space cooling [Davis and Lister, 1974], q = C/ Age,
−2
1/2
with a value of C ∼ 450 mW m /Ma yields a
thermal age of 10 Ma.
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Figure 9. Thermal model for NanTroSEIZE transect. (a) Heat flow values and model profiles for the two models
considered. (b and c) Finite element model showing units. Yellow corresponds to subducting sediment, green corresponds to accreted sediment, and brown corresponds to oceanic crust. Black lines shows plate boundary (lower) and
splay fault (upper). White lines show finite element model structure. Isotherms corresponding to models B and A are
shown in Figures 9b and 9c, respectively.
[39] Our model focuses on the shallow portion of
the subduction zone where we can best estimate its
thermal state, and we explore two models that
bracket the thermal state of the Philippine Sea plate
offshore the Kii Peninsula. In both models we
assume that mantle dynamics have reset the geotherm of the back‐arc crust to a maximum thermal
age. We depress the geotherm through the sedimentary section by adding sediments. Once the
crust subducts it is out of the back arc and ages
relative to the reset thermal age. With this
assumption, steady state thermal models of sub-
duction can be employed, at least for the shallow
portion of the subduction zone. Typical conductive
time constants are calculated from the expression
exp(−z/4at), where z is depth, a is thermal diffusivity (32 km2/Myr), and t is time. If the modeled
configuration has been approximately steady for
the past 5 Ma then temperatures in the prism to a
depth of 25 km are ∼83% of equilibrium. In the
first model (model A) we use the age of the presently subducting oceanic crust so that the thermal
regime of the crust at the Nankai Trough is equivalent to 20 Ma oceanic lithosphere. In model B the
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age of the oceanic crust at the Nankai Trough is
equivalent to 10 Ma oceanic lithosphere. Results
are shown in Figure 9. Model A is relatively cool
and fits the lowest values of heat flow. In contrast,
model B is warmer and fits the heat flow at C0007
and the trend in probe values between 10 and
20 km landward of the trench. The model is only
marginally warmer than the borehole heat flow
on the upper accretionary prism. In model A the
subduction thrust intersects the 100° and 150°C
isotherms at approximately 22 and 48 km, respectively. In model B the subduction thrust intersects
the 100° and 150°C isotherm at approximately
15 and 25 km landward of the deformation front,
respectively. Thus both models predict dehydration
reactions landward of the megasplay fault zone.
[40] These two models have implications for pre-
10.1029/2011GC003593
heat flow varies from 70 mW m−2 near the toe of
the deformation front to 57 mW m−2 at the Kumano
Basin.
[44] 3. Heat flow measurements from IODP sites
are less than seafloor values of heat flow in the
area of the megasplay. This discrepancy likely
reflects the influence of bottom water temperature
variations.
[45] 4. Steady state thermal models of subduction
are consistent with the heat flow data when an
effective age of 10 Ma for the incoming Philippine
Sea plate is used.
Acknowledgments
dicted temperatures along the megasplay. Mechanical models have been used to argue that the
megasplay fault is the primary coseismic plate
boundary near the updip extent of slip [Wang and
Hu, 2006]. The splay fault branches upward
from the plate boundary where the temperature is
approximately 150° and 200°C in models A and B,
respectively. If the updip extent of seismogenesis is
limited by the 150°C isotherm [Hyndman et al.,
1995], then model A predicts the megasplay fault
is largely aseismic whereas model B is consistent
with seismogenesis on the megasplay fault. Future
ocean drilling in the Kumano Basin will test these
predictions.
[46] The authors thank the ship and drilling personnel, staff,
and technicians aboard the D/V Chikyu for their work. We
thank Christian Hardwick and Thomas Etzel for making
divided bar measurements of thermal conductivity at the
University of Utah. We thank H. Hamamoto for sharing his
manuscript prior to publication and M. Yamano for useful discussions. Comments from two anonymous reviewers and the
Associate Editor greatly improved this study. Funding for
this research was provided to Harris and Spinelli through
the U.S. Science Support Program for IODP (NSF
0652315) that is administered by the Consortium for Ocean
Leadership. Schmidt‐Schierhorn was funded by DFG 946
(Deutsche Forschungsgemeinschaft). This research used samples and data provided by the Integrated Ocean Drilling Program (IODP).
8. Conclusions
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Heat flow along the NanTroSEIZE transect: Results Peninsula, Japan