Download Physical properties of upper oceanic crust

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. B4, 2188, doi:10.1029/2001JB001727, 2003
Physical properties of upper oceanic crust: Ocean Drilling Program
Hole 801C and the waning of hydrothermal circulation
Richard D. Jarrard,1 Lewis J. Abrams,2 Robert Pockalny,3 Roger L. Larson,3
and Tetsuro Hirono4
Received 19 December 2001; revised 4 September 2002; accepted 26 November 2002; published 8 April 2003.
[1] The hydrologic evolution of oceanic crust, from vigorous hydrothermal circulation
in young, permeable volcanic crust to reduced circulation in old, cooler crust, causes a
corresponding evolution of geophysical properties. Ocean Drilling Program (ODP) Hole
801C, which obtained the world’s oldest section of in situ, normal oceanic crust,
provides the opportunity to examine relationships among hydrologic properties (porosity,
permeability, fluid flow), crustal alteration, and geophysical properties, at both core plug
and downhole log scales. Within these upper crustal basalts, fluid flux in zones with
high porosity and associated high permeability fosters alteration, particularly hydration.
Consequently, porosity is correlated with both permeability and a variety of hydration
indicators. Porosity-dependent alteration is also seen at the log scale: potassium
enrichment is strongly proportional to porosity. We extend the crustal alteration patterns
observed at Hole 801C to a global examination of how physical properties of upper
oceanic crust change as a function of age based on global data sets of Deep Sea Drilling
Project and ODP core physical properties and downhole logs. Increasing crustal age
entails macroporosity reduction and large-scale velocity increase, despite intergranular
velocity decrease with little microporosity change. The changes in macroporosity and
velocity are significant for pillows but minor for flows. Matrix densities provide the
strongest demonstration of systematic age-dependent alteration. On the basis of observed
decreases in matrix density that are proportional to the logarithm of age, approximately
half of all intergranular-scale crustal alteration occurs after the first 10–15 Myr.
Apparently, crustal alteration continues, at a decreasing rate, throughout the lifetime of
INDEX TERMS: 3015 Marine Geology and Geophysics: Heat flow (benthic) and
oceanic crust.
hydrothermal processes; 7220 Seismology: Oceanic crust; 3035 Marine Geology and Geophysics: Midocean
ridge processes; 5102 Physical Properties of Rocks: Acoustic properties; 8135 Tectonophysics: Evolution of
the Earth: Hydrothermal systems (8424); KEYWORDS: Ocean Drilling Program, basalt, oceanic crust,
logging
Citation: Jarrard, R. D., L. J. Abrams, R. Pockalny, R. L. Larson, and T. Hirono, Physical properties of upper oceanic crust: Ocean
Drilling Program Hole 801C and the waning of hydrothermal circulation, J. Geophys. Res., 108(B4), 2188,
doi:10.1029/2001JB001727, 2003.
1. Introduction
[2] An elegant conceptual model describes the hydrologic
and geophysical evolution of oceanic crust. Hydrothermal
circulation is vigorous in newly created oceanic crust,
resulting in black smokers on the seafloor and alteration
of the upper kilometer of crust. Alteration minerals fill
1
Department of Geology and Geophysics, University of Utah, Salt Lake
City, Utah, USA.
2
University of North Carolina at Wilmington, Center for Marine
Science, Wilmington, North Carolina, USA.
3
Graduate School of Oceanography, University of Rhode Island,
Narragansett, Rhode Island, USA.
4
Department of Earth and Planetary Sciences, Tokyo Institute of
Technology, Maguro, Japan.
Copyright 2003 by the American Geophysical Union.
0148-0227/03/2001JB001727$09.00
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macroporosity (cracks and interpillow voids), causing
increased velocities and densities, decreased porosity and
permeability, and oxidation of magnetic minerals [e.g.,
Jacobson, 1992]. Hydrothermal circulation on ridge flanks
is damped by a relatively impermeable sediment cover
[Anderson and Hobart, 1976], decreased crustal permeability, and decreased thermal buoyancy forces [Anderson et al.,
1977]. Nevertheless, circulation on flanks is responsible for
70% of the advective heat loss from oceanic crust [Stein
and Stein, 1994] and substantial geochemical exchange with
seawater [Mottl and Wheat, 1994; Elderfield and Schultz,
1996].
[3] To decipher the interactions among hydrogeology,
structure, and physical properties in the aging of oceanic
crust, we must examine at least four typical end-member
situations: young crust created at slow and fast spreading
rates and old crust created at slow and fast spreading rates.
5-1
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5-2
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
Figure 1. Locations of DSDP and ODP sites with significant penetrations into normal oceanic crust and
with either downhole logging (Table 2) or core physical properties (Table 3).
The first three of these situations were partly addressed at
Holes 395A (young and slow), 504B (young and moderate
to fast), and 418A (old and slow). Legs 129 and 185 drilled
the fourth end-member: Hole 801C, located in the Pigafetta
Basin at 1838.50N, 15621.60E, northwest Pacific Ocean
(Figure 1). Hole 801C contained the oldest in situ oceanic
crust that has been sampled to date: 166.8 ± 4.5 Ma
[Pringle, 1992]; its biostratigraphic [Bartolini and Larson,
2001] and radiometric basement ages closely match its
predicted age on the basis of extrapolating magnetic anomaly isochrons. This crust was created at a half spreading rate
of 70 – 80 mm/yr [Abrams et al., 1992; Bartolini and
Larson, 2001], close to the fastest rate (90 mm/yr) that is
observed at currently active spreading centers. Leg 185 also
drilled a second example of this fourth end-member: Site
1149 (Figure 1). Unfortunately, basement at Site 1149 could
not be logged, but its core-based geophysical data are
analyzed in this paper.
[4] In the first half of this paper, we analyze physical
properties, both core and downhole logs, of the tholeiites at
Hole 801C, in order to complete the characterization of
oceanic crustal end-members. These specific questions are
addressed for Hole 801C:
1. What is the porosity structure (intergranular, intraflow,
and interflow) of the upper crust, and how is it related to
alteration?
2. What is the relationship of alteration to the velocity
structure of old crust of fast spreading origin? Are the
observations consistent with predictions of the standard
model of geophysical aging of oceanic crust?
3. Do the physical properties of these basalts provide
clues to their hydrothermal circulation history?
[5] In the second half of this paper, we extend the crustal
alteration patterns observed at Hole 801C to a global
examination of how the physical properties of upper oceanic
crust change as a function of age. Basalt coring and logging
by Deep Sea Drilling Project (DSDP) and Ocean Drilling
Program (ODP) provide abundant information on physical
properties of upper oceanic crust, despite generally low
basement core recoveries and rare basement logging [Goldberg, 1997]. These data, acquired and published over a
period of three decades, have not been combined in a
systematic study. Johnson and Semyan [1994] have com-
piled and analyzed physical properties of basalt cores, and
two studies [Carlson and Herrick, 1990; Moos et al., 1990]
have compared results from the three deepest-penetration
sites (Holes 395A, 418A, and 504B). This paper focuses
only on the top 300 m of basalt, thereby permitting us to
document and compare temporal changes among 4 times as
many logged sites (Figure 1). Many basalt sites have cores
but not logs, so some of our additional indicators of geophysical aging of oceanic crust are entirely core based.
2. Data
2.1. Site 801
2.1.1. Coring and Logging
[6] In contrast to all previous sites drilled in the northwest
Pacific, Site 801 succeeded in penetrating mid-Cretaceous
volcanics and obtaining a first-time view of the remnant
Jurassic Pacific plate. Beneath a 192-m-thick section of
Middle Cretaceous volcaniclastics and beneath an apparent
unconformity lies an uninterrupted section of Valanginian to
Bathonian (130– 165 Ma) deep-sea pelagic sediments. The
basal Bathonian sediments are interbedded with the uppermost basalts. Seismic stratigraphy of this portion of the
Pigafetta Basin indicates that Site 801 is located on normal
oceanic crust, free from intersecting faults [Abrams et al.,
1992, 1993].
[7] During ODP Leg 129, Holes 801B and 801C penetrated the upper 131 m of Jurassic basement, with 20 m of
overlap between the two holes [Shipboard Scientific Party,
1990c]. Most of these basalts are thin flows and some are
pillows. A fossilized iron-rich hydrothermal sediment
deposit was found 60 m below the uppermost basalts.
Above this hydrothermal deposit are alkalic basalts, which
are probably sills [Shipboard Scientific Party, 2000a; J. C.
Alt, personal communication, 2002] and are 10 Myr
younger [Pringle, 1992] than the mid-ocean ridge basalt
(MORB) tholeiites from below the deposit [Floyd and
Castillo, 1992]. All recovered samples are affected by
low-grade and highly variable, anoxic submarine alteration,
the intensity of which appears to be unrelated to increasing
depth or lithology [Alt et al., 1992]. The most highly altered
basalts are directly above and below the hydrothermal
sediment deposit. The cooling unit just below the hydro-
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
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5-3
Figure 2. Edited density, velocity, and apparent formation factor logs from Hole 801C tholeiites,
compared with core plug measurements. Note the consistency of core and log measurements for velocity
and density, but not for apparent formation factor. Extrusive units are from Shipboard Scientific Party
[2000a].
thermal deposit is highly fractured and extremely altered,
with a late, oxidative type of alteration that is suggestive of
proximity to a hydrothermal system [Alt et al., 1992].
[8] Although Hole 801C was drilled during Leg 129, no
downhole experiments were undertaken in the hole at that
time. ODP Leg 144 returned to this site and completed both
packer [Larson et al., 1993] and logging [Jarrard et al.,
1995] measurements. A complete suite of in situ geophysical logs was obtained through the upper 100 m of basalt,
including velocity, density, neutron porosity, resistivity,
three-component magnetometer, geochemistry, and Formation MicroScanner1 logs. ODP Leg 185 returned to this
hole and deepened it by 339.3 m for a total of 474 m
penetration onto oceanic crust [Shipboard Scientific Party,
2000a]. Recovered cores consisted almost entirely of
MORB tholeiites, both flows and pillows, with rare hyaloclastites and sediments plus a second hydrothermal unit
(Figure 2). Downhole logging extended from a bridge at
70– 90 m above total depth, up to the bottom of casing at
483 m below seafloor (mbsf). This includes the zone
previously logged on Leg 144. Nearly the same suite of
logs was run as on Leg 144, except that geochemical
logging was confined to the spectral gamma tool. Resistivity logging on Leg 185 used the dual laterolog rather than
dual induction tool.
[9] Physical properties of the upper hydrothermal zone and
overlying alkali basalts have been presented and analyzed
previously [Busch et al., 1992; Wallick et al., 1992; Larson et
al., 1993; Jarrard et al., 1995] and are not presented here.
Our analyses focus on the tholeiitic section, first encountered
on Leg 129 but mostly cored and logged during Leg 185.
2.1.2. Core Measurements of Physical Properties
[10] Busch et al. [1992] and Wallick et al. [1992] measured core-based physical properties of upper Hole 801C
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5-4
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
basalts. These measurements included index properties
(bulk density, matrix (or grain) density, porosity) and P
wave velocity at atmospheric pressure. Wallick et al. [1992]
also measured velocity versus pressure. Our analyses
exclude the data of Wallick et al. [1992], because of
incomplete removal of interstitial water during drying
[Jarrard et al., 1995]. Leg 185 basalt physical properties
sampling used a much closer sample spacing than at any
previous deep crustal site: approximately three samples per
core. At each horizon, a wedge-shaped sample was taken for
index properties [Shipboard Scientific Party, 2000a] and an
adjacent cube was cut for three-axis measurement of velocity [Shipboard Scientific Party, 2000a], resistivity and X-ray
computed tomography (CT) imaging [Hirono and Abrams,
2002]. An additional 235 measurements of horizontal P
wave velocity were made on split cores.
[11] Alteration of the Leg 185 basalt cubes was estimated
with light absorption spectroscopy (LAS). LAS measures
the spectrum of light reflected from a rock surface. Light is
absorbed by minerals at and near the rock surface due to
both electronic and vibrational processes [Clark et al., 1990;
Clark, 1995]. Types of identifiable minerals depend on the
frequency band of the instrument. We use the FieldSpec Pro
FR Portable Spectroradiometer1 because of its wide bandwidth (350 – 2500 nm). At near-infrared wavelengths (950 –
2500 nm), normal modes of characteristic vibrations of OH
bonds occur; diagnostic absorption bands for water, MgOH, Al-OH, and Fe-OH are evident.
[12] LAS has not previously been used as a technique for
basalt analysis, yet it has several advantages over traditional
DSDP/ODP analytical techniques: measurements take only
seconds and are nondestructive, and LAS is particularly
sensitive to hydration and smectite concentration, both of
which have presented challenges to interlaboratory measurement consistency. The volumetrically dominant minerals
in these basalts are pyroxene and plagioclase. Plagioclase is
spectrally featureless and therefore undetectable by LAS.
The dominant spectral signature in these basalts is from
pigeonite, the pyroxene. Montmorillonites, with an OH
absorption band at 1400 nm and a strong water absorption
band at 1930 nm, can be detected at concentrations of only a
few percent [Vanden Berg and Jarrard, 2002]. Trough
depths at these two wavelengths, each normalized to adjacent wavelengths outside the absorption band, are highly
correlated (R = 0.882), so we combine them into a single
measure of relative abundance of montmorillonite among
the Leg 185 basalts. Pore water has the same OH and water
absorption bands as hydration minerals, so it is important to
use LAS only on thoroughly dried samples. To assure
complete drying, we evacuated the samples for 24 hours
at a vacuum pressure of 9 – 11 Pa.
[13] Permeabilities of 13 of these cube samples were
measured by Terra Tek1, a commercial petrophysical
analysis laboratory. Initial reconnaissance testing consisted
of measurements of eight relatively porous samples using a
low-pressure (300 psi or 2 MPa), steady state air flow
technique on jacketed samples. Unfortunately, five of the
eight samples were below the technique’s lower limit of
1017 m2 (0.01 mdarcy) (Table 1). Furthermore, of the three
that were measurable, one developed a crack during testing
and is therefore nonrepresentative of intergranular permeability, and one has a permeability that is barely resolvable
Table 1. Permeability Measurements on Cores From Hole 801C
Permeability, m2
Depth, mbsf
606.12
607.22
607.74
615.11a
616.91
620.12
673.35
693.07
720.68
757.00
815.17
900.32
901.07
Steady State
17
<10
<1017
2.3 1016
5.3 1016
<1017
<1017
1.2 1017
<1017
Pulse Decay
19
8.25 10
1.38 1018
4.04
6.66
1.24
5.22
1019
1019
1019
1019
3.18 1017
2.46 1019
Porosity, %
18.0
13.4
15.6
7.5
8.6
10.4
7.6
6.1
10.5
11.0
5.9
8.1
7.7
a
Sample cracked during measurement.
by this technique. Our second stage of testing measured
permeability of eight samples with a pulse decay technique,
capable of accurate permeability measurements at levels as
low as 1020 m2. At these very low permeabilities, even
traces of residual pore water or surface humidity may bias
measurements, so care was taken to minimize both.
2.1.3. Downhole Logs
[14] Five downhole logs of the Hole 801C tholeiitic
interval were analyzed: density, sonic (velocity), resistivity,
neutron, and potassium.
[15] Density logging uses a pad-type tool that requires a
good contact with the borehole wall for reliable performance. Most previous logging of oceanic crust used density
tools that frequently lost pad contact with the borehole wall;
consequently, these density logs were of generally poor
quality [Carlson and Herrick, 1990]. Beginning on ODP
Leg 125 the hostile environment lithodensity tool was used.
Use of this tool at Hole 801C resulted in a density log that
appears to be of generally fair-good quality. The caliper log
indicates that hole size is nearly uniform throughout most of
the hole, except for poor hole conditions in the interval
627– 714 mbsf, a unit consisting of thin flows and pillows
(Figure 2). Loss of pad contact and resulting low-density
readings are most common in this interval. We used a
combination of caliper, r, velocity, and photoelectric
factor logs to identify zones with loss of pad contact and
to guide deletion of associated unreliable density data.
Approximately 21% of the data were deleted in this manner.
As discussed in a subsequent section, however, interlog
comparisons suggest that spurious values are common even
in this edited density log. Figure 2 shows the edited density
log; the unedited log is shown by Shipboard Scientific Party
[2000a].
[16] The basalt interval at Hole 801C was logged with the
borehole-compensated sonic tool, which generates four
source-receiver travel times at each measurement point. A
few of the redundant measurements at each depth were
obviously unreliable due to cycle skips and ‘‘road noise’’
(reverberations due to tool drag over rough borehole wall).
We employed a simple reprocessing algorithm [Shipboard
Scientific Party, 1987] that rejected the unreliable data,
permitting calculation of a more reliable velocity log. The
reprocessed velocity log of Figure 2 appears to be of good
quality.
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
[17] Leg 185 resistivity logging of Hole 801C [Shipboard
Scientific Party, 2000a] used the dual laterolog, which has
two depths of penetration (shallow and deep). This tool
provides superior accuracy to the dual induction tool (used
on Leg 144) in high-resistivity formations such as some
basalts, because dual-induction response saturates at resistivities near 2000 m. In general, resistivity logs have the
highest signal-to-noise ratio of any log obtained in ODP.
The shallower-penetration logs are, however, sensitive to
changes in hole size: In enlarged portions of the hole, they
underestimate formation resistivity. Consequently, we confine our analyses to the deep-resistivity log, which has a
depth of penetration so great that it is insensitive to hole size
variations.
[18] Rock resistivity depends mainly on resistivity of the
formation fluid, which is a function of both salinity and
temperature. Apparent formation factor, the ratio of rock
resistivity to fluid resistivity, removes this effect. We converted core-based resistivities to apparent formation factors,
assuming a fluid resistivity of 0.2 m (seawater salinity
and room temperature of 21C). Assuming seawater salinity
and the temperatures measured within the borehole at the
time of logging [Shipboard Scientific Party, 2000a], log
resistivities were converted to apparent formation factors
(Figure 2). Because the time between logging and cessation
of drilling is small compared to the time spent drilling and
circulating drilling fluids, we suspect that near-borehole
temperatures are similar to those measured within the
borehole. If the zone reached by the deep resistivity tool
attained equilibrium temperature by the time of logging,
then temperatures are 10 higher than we assumed and
formation factors are 25% lower than we calculated. This
uncertainty is relatively small, compared to the 2 orders of
magnitude variations in formation factor observed within
the logged interval.
[19] An alternative approach to Leg 185 formation temperatures is to compare the upper part of the Leg 185
resistivity log to the log obtained on Leg 144. If the dual
laterolog and induction tools have identical log responses,
then the ratio of these logs can be converted to a log of
temperature change. We found, however, that the differences between these two logs were too great to be attributable to temperature change; tool response is probably
responsible.
[20] ODP neutron porosity logs consistently overestimate
the porosity of basalts [Anderson et al., 1985; Broglia and
Moos, 1988; Lysne, 1989; Moos, 1990]. Broglia and Ellis
[1990] utilized logs from five sites to analyze the relative
contributions of several factors to this systematic bias. They
concluded that the most important factor is the ODP
technique of running the neutron porosity tool without
appropriate eccentralization. At the one site (Site 735) that
was logged with a bowspring-eccentralized neutron tool,
this bias was absent. Although neutron tools are ‘‘compensated’’ in the sense of applying some correction for tool
standoff from the borehole wall, Broglia and Ellis [1990]
found that this compensation is insufficient for correction of
neutron logs that lack eccentralization.
[21] The Hole 801C neutron logs run on Legs 144 and
185 were corrected for standoff. Jarrard et al. [1995]
compared Leg 144 density-based porosity to neutron porosity and found that the neutron tool overestimated porosity
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5-5
by 5– 25%. A similar cross plot for the Leg 185 logs (not
shown) exhibits a nearly identical pattern to that for Leg
144. This discrepancy, which is higher than those identified
by Broglia and Ellis [1990] in other ODP crustal sites,
probably results from a combination of interlayer water in
alteration minerals and calibration to limestone matrix
density rather than to basalt matrix density. Broglia and
Ellis [1990] corrected neutron logs from several crustal sites
for both effects. For Hole 801C, however, too few H2O+
analyses are available for alteration correction, which compromises the usefulness of the Hole 801C neutron logs.
2.2. Other Sites
[22] From the DSDP and ODP Initial Reports and their
associated CD-ROM databases, we have extracted four
types of data: velocity logs, core plug velocities, core index
properties (bulk density, matrix density, and porosity), and
core-based identification of volcanic style.
[23] From 1200 sites drilled by DSDP and ODP, we
selected those that fulfill the following criteria: (1) velocity
logging of at least 40 m of basement and (2) basement
composed of normal oceanic crust formed in an open ocean
environment. This second criterion excluded sites on crust
formed by back arc spreading, oceanic plateau or seamount
volcanism, or at a heavily sedimented spreading center; all
three environments are likely to differ from normal oceanic
crust in both volcanic style (especially proportion of intrusives versus extrusives) and in hydrothermal alteration.
Thirteen holes met both criteria: 395A, 396B, 417D,
418A, 504B, 556, 558, 564, 765D, 770C, 801C, 843B,
and 896A (Table 2 and Figure 1). Basement logs for these
sites [Kirkpatrick, 1978; Salisbury et al., 1979; Cann and
Von Herzen, 1983; Hill and Cande, 1985; Broglia and
Moos, 1988; Moos, 1990; Jarrard and Broglia, 1991;
Goldberg and Moos, 1992; Jarrard et al., 1995; Shipboard
Scientific Party, 1978a, 1978b, 1979a, 1979b, 1983, 1985a,
1985b, 1985d, 1990a, 1990b, 1990c, 1993, 1998, 2000a]
vary in length, from a minimum of 41 m at Site 843 to 1827
m at Hole 504B. Most sites have 90– 200 m of velocity log
(Figure 3), and only three (395A, 418A, and 504B) have
more than 300 m. We confined our analyses to the top 300
m of basement, thereby preventing our intersite comparisons from being biased by intrasite velocity gradients.
Except for a possible velocity gradient within the top 30
m of basement, no systematic depth-dependent changes in
velocity are evident within the top 300 m of basement
(Figure 3).
[24] Most DSDP and ODP velocity logging has utilized
the Schlumberger borehole-compensated sonic tool, which
determines P wave travel time by a first-break thresholding
criterion: the first energy exceeding a preset threshold is
assumed to be that of the initial P wave peak [e.g., Serra,
1984]. Raw velocity logs for basalts generally have many
unreliable intervals, because cycle skips and ‘‘road noise’’
(reverberations due to tool drag over rough borehole wall)
induce failure of the threshold technique. For most sites, we
employed a simple reprocessing algorithm [Shipboard Scientific Party, 1987] that we have developed to take advantage of redundancy of measurements, reject the unreliable
data, and use retained data to calculate a more reliable
velocity log. Original first-break travel time logs were not
available for Sites 396 and 417, so their velocity logs could
5-6
0.011
0.017
0.010
0.012
0.003
±
±
±
±
±
0.56 ± 0.015 0.54
0.70 ± 0.019 0.39
0.29 ± 0.018 0.28
0.40
0.17 ± 0.003 0.16
0.017
0.040
0.012
0.012
0.005
±
±
±
±
±
0.54
0.70
0.24
0.40
0.15
CL
K, %
CL
K, %
CL
K, %
CL
0.21 ± 0.036 0.37 ± 0.019 0.33 ± 0.016
0.12 ± 0.007 0.19 ± 0.003 0.17 ± 0.003
a
± 0.03
4.67
5.61
4.62
5.55
5.17
5.47
5.20
4.83
5.68
5.17
4.67
5.22
4.77
4.73
±
±
±
±
±
1.9
2.4
2.8
2.2
7.0
5.7
2.0
1.1
1.9
3.3
± 3.2
± 8.6
± 0.9
0.4
2.2
1.5
92 – 393
214 – 336
343 – 448
324 – 624
287 – 574
462 – 560
409 – 543
284 – 336
926 – 1153
427 – 519
530 – 820
257 – 297
190 – 403
92.0
150.5
343
324
274
461.5
408
284
921
423.2
530
242.5
179
17
20
15
15
38
10
11
10
37
45
90
?
38
7.0
9.4
108
110
5.9
32
35.5
34
128
42
167
110
5.9
Average macroporosity (f), velocity (V ), and potassium (K), along with 95% confidence limits, are shown. CL, confidence limits.
4.71
4.27
5.02
4.93
4.70
5.04
4.36
5.18
5.08
4.53
5.01
4.77
4.67
0.03
0.07
0.04
0.05
0.04
0.03
0.05
0.03
0.03
0.06
0.05
CL
±
±
±
±
±
±
±
±
±
±
±
4.55
4.25
4.90
4.87
4.56
5.04
4.75
5.14
4.99
4.44
4.87
V, km/s
CL
CL
f,%
1.8 6.1 ± 1.8
3.0 10.1 ± 2.8
2.5 3.3 ± 2.1
1.4 3.5 ± 1.2
2.8 8.1 ± 2.3
3.1 3.5 ± 2.9
7.2 5.3 ± 7.0
1.6 2.8 ± 1.7
3.1 0.5 ± 3.0
2.8 3.7 ± 1.7
2.3 1.7 ± 0.9
2.2 ± 1.9
11.2 ± 5.6 8.1 ± 2.8
8.4
10.1
4.1
4.4
10.1
2.6
4.4
3.2
2.4
4.3
1.1
CL
f, %
CL
Site
Age, Spreading Rate, Basement Depth, Log Interval
Ma
mm/yr
mbsf
Analyzed, mbsf f, %
0.9 ± 3.1
±
±
±
±
±
±
±
±
±
±
±
V, km/s
±
±
±
±
±
±
±
±
±
±
±
±
±
0.07
0.22
0.05
0.09
0.05
0.16
0.20
0.10
0.05
0.08
0.04
0.07
0.05
V, km/s
±
±
±
±
±
±
±
±
±
±
±
±
±
0.03
0.07
0.04
0.04
0.04
0.03
0.06
0.03
0.03
0.04
0.03
0.07
0.03
All
Pillows
Potassium
Flows
All
Pillows
Velocity
Flows
All
Macroporosity
Pillows
Flows
Table 2. Sites With Significant Downhole Logging of Normal Oceanic Crusta
0.22 ± 0.008 0.40 ± 0.006 0.35 ± 0.007
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
395A
396B
417D
418A
504B
556
558
564
765D
770C
801C
843B
896A
EPM
not be reprocessed. For Sites 418 and 504, we used velocity
logs calculated by semblance analysis of waveforms [Moos
et al., 1990]. Velocity and resistivity are both controlled
primarily by porosity, but resistivity measurements are less
affected by poor borehole conditions. We removed velocity
values in intervals where velocity spikes to low values were
inconsistent with resistivity character.
[25] Core velocity measurements used here (Figure 3)
were shipboard measurements, made at atmospheric pressure and reported by the shipboard scientific parties. For
some sites, postcruise studies determined velocity as a
function of pressure, but we have chosen not to mix the
different data types.
[26] Volcanic style is evaluated from core descriptions by
Shipboard Scientific Party [1978a, 1978b, 1979a, 1979b,
1983, 1985a, 1985b, 1985d, 1990a, 1990b, 1990c, 1992,
1993, 2000a]. When core recovery is high, distinguishing
pillow basalts from sheet flows is relatively straightforward.
Often, however, low core recovery lends ambiguity to this
discrimination. For example, Hole 843B yielded only lowrecovery spot cores that were all described as flows, though
Formation MicroScanner1 logs show that some pillows are
present [Goldberg and Moos, 1992]. If core recovery is
sufficient to discriminate flows from pillows, then the
continuous records provided by logs usually permit a
refinement of unit boundaries because geophysical properties generally change much more at the boundaries of such
units than within them. These slight revisions of boundaries,
either shipboard or shore-based [e.g., Broglia and Ellis,
1990], are not always applied, so we have redone them for
all logged sites. Similarly, we have confirmed or modified
the log depth shifts for all sites; in some cases, cable stretch
has necessitated up to several meters constant depth shift of
the ‘‘final’’ logs available from CD-ROM and the LamontDoherty Earth Observatory log data repository. Although
nearly all penetrated basement rocks are either pillows or
massive sheet flows, several other kinds of igneous rocks
are encountered occasionally. Breccias are the most common of these; we group breccias with pillows, because the
boundary between the two is gradational in morphology and
probably also in extrusive intensity. At the shallow subbasement depths considered here, sills are rare. Recovery of
sediment interlayers is usually less than a few centimeters,
partly because of core recovery bias. Each of the other
encountered rock types is present at only one site: basaltic
sand and gravel at Hole 396B, serpentinized gabbro at Site
556, serpentinite at Site 558, and hydrothermal deposits at
Hole 801C.
[27] A second search of the DSDP and ODP databases
and publications (Initial Reports and Scientific Results)
focused on index properties of basalts (Table 3). Several
DSDP sites have measurements of basalt bulk density but
not porosity or matrix density; results from these are
tabulated by Johnson and Semyan [1994] but excluded
from our analyses. Again, only sites on normal oceanic
crust were considered; data from sedimented ridges (e.g.,
Juan de Fuca and Gulf of California) were excluded. Data
from multiple holes at the same site were combined, except
for two sites (417A&D, 1149B&D) at which the holes were
widely spaced or clearly in contrasting hydrothermal settings (e.g., altered basement high versus less altered basement low). Useful index data are available for 25 sites
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
EPM
5-7
Figure 3. P wave velocity logs, reprocessed and edited, and measurements of core plug velocity, for the
13 DSDP and ODP sites of this study. Sites are plotted in order from youngest (504B) to oldest (801C), to
graphically demonstrate the general pattern of increasing agreement of the two types of data with
increasing age. The horizontal scale on all plots is velocity, ranging from 3 km/s at each left margin to 7
km/s at each right margin. All plots extend to 300 m subbasement, the maximum depth of these analyses.
(Table 3 and Figure 1), almost double the number of sites
with velocity logs (Table 2).
3. Site 801 Analysis
3.1. Permeability
[28] Permeability is the architect of the crust’s hydrothermal systems. Nearly all measurements of the permeability of upper oceanic crust are averages for portions of a
borehole tens to hundreds of meters in extent. Fisher [1998]
provides an excellent review and synthesis of these bulk
permeability measurements and their implications. This
large-scale crustal permeability is a critical control on
hydrothermal circulation patterns and associated heat flux
because the volumetrically dominant flux is channelized
within the most permeable zones, particularly large open
fractures [Fisher, 1998; Fisher and Becker, 2000].
[29] Bulk permeabilities for the top 1200 m of oceanic
crust suggest two main layers: the top 500 m has permeabilities of 10141013 m2, and the next 700 m has
permeabilities of 1017 m2 [Fisher, 1998]. These bulk
permeabilities are all from <8 Ma crust. A single measurement from older crust is that at Hole 801C, where testing of
a 93-m interval indicated a bulk permeability of 8 1014 m2; more likely, flow is confined to an 18-m hydrothermal zone at the top of the tholeiites, and bulk permeability is accordingly higher, 4 1013 m2 [Larson et al.,
1993]. Drilling-induced temperature anomalies measured on
5-8
EPM
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
Table 3. Average Index Properties of Basalts for Sites on Normal Oceanic Crust, With 95% Confidence Limits
Site
Hole
Age, Ma
332
333
334
395
396
409
410
412
417
417
418
470
504
505
543
556
559
562
564
597
648
765
770
801
896
1149
1149
A&B
A
3.5
3.5
8.9
7.0
9.4
2.4
9
1.6
108
108
110
15
5.9
4
80
32
32
16
34
28.5
0
128
42
167
5.9
132
132
&A
B
&A
A
A
D
A
A
B
B
A
C
B
C&D
B&C
C
A
B
D
Matrix Density
f > 5%, g/cm3
95% CL
2.963
2.974
± 0.015
± 0.015
2.954
2.847
2.970
2.903
3.010
2.835
2.924
2.878
2.940
2.981
2.988
2.893
2.955
2.947
2.970
3.005
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
0.018
0.086
0.022
0.035
0.013
0.028
0.018
0.017
0.050
0.007
0.033
0.100
0.040
0.092
0.034
0.024
2.999
2.835
2.839
2.880
2.936
2.796
2.829
±
±
±
±
±
±
±
0.041
0.031
0.042
0.021
0.016
0.074
0.027
Matrix Density,
g/cm3
2.942
2.974
2.992
2.955
2.855
2.970
2.903
3.005
2.852
2.925
2.899
2.958
2.987
2.986
2.898
2.947
2.938
2.960
2.995
3.103
2.998
2.867
2.845
2.900
2.948
2.796
2.832
Leg 185 clearly delineate this high-permeability zone, as
well as a thinner permeable zone at 705– 712 mbsf [Shipboard Scientific Party, 2000a]. This deeper zone is very
porous, with the lowest formation factors and velocities of
the entire tholeiitic interval (Figure 2).
[30] Bulk permeability measurements cannot address the
problem of intergranular-scale permeability and the associated fluid flow that is responsible for pervasive intergranular-scale crustal alteration. Intergranular permeabilities of
most rocks are controlled by pore geometry, which in turn is
dependent on porosity and lithology. This dependence
commonly results in a linear relationship between porosity
and ln(permeability) for all rocks from the same formation
[Nelson, 1994]. Very few core plug measurements of basalt
permeability have been undertaken. Young basalts from
DSDP Leg 70 (0.5 –2.7 Ma) and Hole 504B (5.9 Ma) have
permeabilities of mostly 1020 to 1018 m2 that are weakly
dependent on porosity [Karato, 1983a, 1983b]. Permeabilities of 110 Ma crust at Holes 417D and 418A are generally
1021 to 1017 m2 [Johnson, 1979a; Hamano, 1979] and
moderately correlated with porosity. These intergranular
permeabilities are 3 – 8 orders of magnitude lower than
bulk permeabilities for the upper layer, which in turn are
less than regional-scale permeabilities [Fisher et al., 1994;
Fisher and Becker, 2000]. Other basalt permeability measurements may not be representative of open ocean crust:
Juan de Fuca sedimented ridge [Christensen and Ramananantoandro, 1988], Tonga-Kermadec [Christensen and
Ramananantoandro, 1988], and Arctic Ocean [Aksyuk et
al., 1992].
[31] Our samples from Hole 801C (Table 1), along with
mostly lower-porosity samples from other sites on normal
oceanic crust [Johnson, 1979a; Hamano, 1979; Karato,
1983a, 1983b], exhibit a linear relation between porosity
and ln(permeability) (Figure 4a). Compared to typical
95% CL
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
0.016
0.015
0.032
0.010
0.030
0.022
0.035
0.011
0.026
0.013
0.016
0.019
0.005
0.023
0.052
0.027
0.080
0.026
0.023
0.019
0.015
0.016
0.025
0.016
0.009
0.074
0.026
Porosity, %
7.36
11.13
4.75
5.27
5.51
16.27
23.43
5.31
8.50
6.67
8.36
4.71
5.34
6.15
5.77
6.77
12.23
5.05
7.79
3.53
5.03
4.02
5.34
7.56
5.73
10.27
8.27
95% CL
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
1.02
3.72
3.13
0.58
1.33
0.96
2.05
0.76
3.59
0.90
1.70
1.33
0.43
1.48
1.50
2.41
5.06
1.52
2.18
1.60
0.76
0.64
0.79
0.89
1.14
4.69
1.50
Bulk Density,
g/cm3
2.800
2.754
2.897
2.852
2.752
2.650
2.456
2.897
2.698
2.796
2.742
2.865
2.881
2.864
2.788
2.814
2.701
2.863
2.839
3.030
2.897
2.792
2.747
2.759
2.842
2.616
2.682
95% CL
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
0.018
0.085
0.088
0.016
0.039
0.028
0.038
0.013
0.081
0.022
0.042
0.037
0.011
0.038
0.072
0.055
0.128
0.029
0.029
0.033
0.023
0.023
0.031
0.026
0.027
0.134
0.038
trends for sedimentary rocks, however, the basalt scatter is
much higher and permeabilities are lower by 3 orders of
magnitude. Both differences are attributable to the isolation
of most vesicles from available flow paths. The relationship
between permeability and porosity, though noisy, indicates
that porosity can be used as an indirect indicator of permeability and therefore of relative fluid flux. Permeability
controls alteration by determining water-rock ratio. Consequently, measurements of LAS-based alteration index for
Hole 801C are correlated with both permeability (Figure 4b)
and porosity (Figure 5c). On the basis of our small sample
from Hole 801C, alteration does not appear to cause a major
change in the porosity/permeability relation.
[32] On the basis of the freshness of most non-801C
samples in Figure 4a and the regression trend of Figure
4b, the minimum permeability for sufficient water-rock ratio
to generate significant alteration is 1019 m2, which
occurs at a porosity of 5%. Flows, with porosities of
mostly <5%, consequently undergo much less alteration
than more porous pillows. However, these threshold permeability and porosity values are only generalizations,
because water-rock ratio and alteration depend not only
on local intergranular permeability but also on proximity to
high-permeability cracks.
3.2. Alteration
[33] Basalt physical properties at the centimeter scale
measured in core plugs may differ from those at the meter
scale measured by logs, because of differences in type of
porosity and in alteration history. At plug scale, porosity
consists mainly of vesicles and microcracks, whereas logscale porosity also includes fractures. At plug scale, alteration converts glass, olivine, and plagioclase mainly to clay
minerals, thereby decreasing matrix density and matrix
velocity. Porosity may increase due to this hydration, or
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
EPM
5-9
1992; Jarrard et al., 1995]. An association between alteration and porosity was also evident, probably because of
two factors: (1) alteration increases porosity and (2) high
porosity promotes alteration because of high permeability
and therefore enhanced fluid flow. Figure 5b confirms that
matrix densities of Hole 801C tholeiites decrease with
increasing porosity, although the relationship is weaker than
was observed for the short tholeiite interval cored during
Leg 129.
[35] CHN-based hydration measurements were not available for the Leg 185 index property samples, so we
determined relative variations in hydration with LAS. Figure 5a shows that matrix density for these basalts is only
subtly correlated with LAS-based hydration, in contrast to
the much better association between matrix density and
hydration for the shorter interval analyzed by Jarrard et al.
[1995]. Increasing hydration is associated with increasing
porosity (Figure 5c), with a higher correlation (R = 0.69)
than observed by Jarrard et al. [1995].
Figure 4. (a) Relationship between permeability and
porosity for core plugs from Hole 801C (Table 1), compared
to published data for young crustal sites (Hole 504B and
Leg 170) [Karato, 1983a, 1983b] and old crustal sites
(Holes 417D and 418A) [Hamano, 1979; Johnson, 1979a].
(b) Relationship between permeability and LAS-based
alteration intensity for core plugs from Hole 801C.
decrease due to precipitation of carbonates or zeolites. At
log scale and larger, porosity is thought to decrease due to
filling of cracks and interpillow voids by alteration minerals, carbonates, quartz/chalcedony, or Fe-oxyhydroxides.
[34] Basalt core descriptions provide only a subjective
description of variations in alteration. The H2O+ of basalts,
as measured by CHN analyzer, does provide a semiquantitative measure of the extent of alteration-induced hydration
[Alt et al., 1992]. Typical H2O+ contents of 3 – 6 wt % for
the upper extrusives contrast with initial MORB H2O+ of
only 0.12 wt % in fluid inclusions [Sobolev and Chaussidon, 1996]. For the portion of Hole 801C cored during Leg
129, the index property measurements of Busch et al.
[1992] were paired with geochemical analyses of Alt et al.
[1992], permitting demonstration of a close correlation
between increasing hydration (greater H2O+) and decreases
in velocity, matrix density, and matrix velocity [Busch et al.,
3.3. Density
[36] Formation bulk density (rb) depends on fractional
porosity (f), matrix (or grain) density (rma), and fluid
density (rf), according to the relationship rb = rma (1 f) +
rf f. Matrix density is the average density of the minerals
forming the solid part of the rock, including any alteration
minerals. This definition and an analogous one for matrix
velocity are universal in physical properties analyses but not
log analyses. Many log interpretations, particularly hydrocarbon-related ones, confine matrix density and matrix
velocity to the ‘‘clean’’ portion of a rock, often with
corrections for any clay content. Fractional porosity is
generally much more variable than matrix density and fluid
density. Thus the density log is a relatively straightforward
porosity log, and both rma and rf often can be assumed to be
constant. For upper crustal basalts, however, matrix density
often decreases with increasing porosity, because of greater
alteration in the more porous and therefore more permeable
rocks. This pattern has been observed in several other
studies [Hamano, 1979; Christensen et al., 1980; Carlson
and Herrick, 1990; Jarrard and Broglia, 1991], including
prior analyses of the upper basalts from Hole 801C [Busch
et al., 1992; Jarrard et al., 1995]. Jarrard et al. [1995]
modified the equation above to include this effect, based on
the upper 801C basalts, and we used their equation to
convert the new Hole 801C density log into a porosity log.
3.4. Velocity
[37] Porosity is the most important variable controlling
both density and velocity in the 801C basalts. This common
control imparts a strong correlation between density and
velocity, routinely noted for individual sites and summarized for core plugs from multiple sites by Busch et al.
[1992]. Figures 6a and 6b compare velocities to densities
for the Hole 801C tholeiites on the basis of core and log
data.
[38] The relation between density and velocity can be
predicted on the basis of a combination of the theoretical
density equation above and the Wyllie et al. [1956] equation: V 1 = (1 f)/Vma + f/Vf, where V is whole rock
compressional wave velocity, f is fractional porosity, Vma is
matrix velocity, and Vf is pore fluid velocity. This equation
EPM
5 - 10
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
Figure 5. Relationships among porosity, alteration (based on light absorption spectroscopy, or LAS),
matrix density, and apparent formation factor, based on core plug measurements for tholeiites from Hole
801C. In general, higher porosities promote hydration reactions generating clays, thereby increasing
LAS-based alteration intensity and lowering matrix density and formation factor. Solid dots: basalts; open
circles: interpillow material.
is only an empirical approximation; theoretical equations
[e.g., Gassmann, 1951; Kuster and Toksöz, 1974] require
elastic moduli that are rarely measured. Matrix velocities for
basalts [e.g., Serra, 1986] vary due to changes in alteration
and composition (particularly proportion of mafic minerals).
The Site 801 basalts cored during Leg 129 have a zeroporosity matrix velocity of 6.5– 6.8 km/s [Jarrard et al.,
1995], whereas those cored during Leg 185 have a matrix
velocity of 6.0– 6.3 km/s.
[39] Assuming a matrix velocity of 6.3 km/s and matrix
density of 3.0 g/cm3, the predicted relationship between
velocity and density at Site 801 is shown by a curve on
Figures 6a and 6b. This curve fits the lowest-porosity core
and log data reasonably well, but it systematically mismatches higher-porosity data. We attribute this mismatch to
porosity-dependent variations in matrix density and matrix
velocity, caused by increased alteration at higher porosities.
Jarrard et al. [1995] noted a similar pattern for core plug
data from upper 801C. They found, in contrast, that a
predicted curve incorporating regression-based porositydependent matrix densities and matrix velocities fit not only
the 801C core plug data but also the synthesis by Busch et
al. [1992] of core plug measurements from various sites.
Figures 6a and 6b confirm that this pattern holds throughout
the deepened tholeiitic interval at 801C. The pattern holds
for both core and log data, despite differences in type of
porosity.
[40] Several other studies have examined the relationship
between basalt porosity and velocity, based on core or log
data. Those that have examined variations in matrix velocity
have noted reduced matrix velocities at high porosity,
attributed to alteration [Carlson and Herrick, 1990; Jarrard
and Broglia, 1991]. A velocity/porosity cross plot for Hole
801C (not shown) indicates that low-alteration samples
(based on LAS) are only subtly higher in velocity, by
200 m/s, than more altered samples.
3.5. Formation Factor
[41] The deep and shallow laterologs differ not only in
depth of penetration but also in overall current geometry:
the shallow measurement path is predominantly vertical and
the deep path is more horizontal. Pezard and Anderson
[1989] and later Pezard et al. [1996] used this distinction to
identify basalt intervals at Hole 504B with dominantly
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
EPM
5 - 11
Figure 6. (left) Core-based and (right) log-based relationships among velocity, bulk density, and
apparent formation factor for Hole 801C tholeiites. Log porosities are calculated from the bulk density
log. Also shown for comparison to the data of Figures 6a and 6b are predicted relationships assuming
either constant matrix values or porosity-dependent matrix values. Note that porosity-dependent matrix
values give the most successful prediction of observed variations. Open circles: basalts; crosses:
interpillow material.
horizontal (Rdeep < Rshallow) or vertical (Rdeep > Rshallow)
fractures, and a similar approach has been used at nearby
Hole 896A [de Larouzière et al., 1996]. Comparison of the
deep and shallow laterologs for Hole 801C tholeiites shows
generally very close agreement [Shipboard Scientific Party,
2000a]. A cross plot, not shown here, indicates a nonlinear
relationship: Shallow resistivity is systematically lower than
deep resistivity at the highest and lowest resistivities. This
behavior implies that differences in tool response (internal
calibrations?) overwhelm any possible influence of resistivity anisotropy on the logs.
[42] Our analyses use formation factor (F = Ro/Rw) rather
than the resistivity log (Ro), thereby removing the influence
of variations in fluid resistivity (Rw). If clays have a minor
EPM
5 - 12
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
contribution to formation factor, then F is related to fractional porosity by the Archie [1942] equation, as generalized
by Winsauer and McCardell [1953]: F = a fm. Surface
conduction on clays can have a significant effect on basalt
resistivity, particularly at elevated temperatures [Olhoeft,
1981]. Consequently, calculations of porosity from resistivity for Hole 504B are based on log-based estimates of both
pore fluid and clay conduction [Pezard, 1990; Pezard et al.,
1996]. At Sites 768 and 770, in contrast, clay conduction
was detectable but small enough to neglect without substantially biasing porosity determination [Jarrard and
Schaar, 1991]. Our analyses for Hole 801C do not include
estimates of clay conduction, so we calculate apparent
formation factors rather than intrinsic, clay-free formation
factors.
[43] Creation of hydrous minerals during alteration is
expected to decrease both apparent formation factor and
matrix density. As previously mentioned, this process is
often most advanced in the highest-porosity, most permeable zones. The combination of high pore fluid conduction
and high clay conduction in porous, altered zones generates
a correlation between apparent formation factor and matrix
density, as seen in Figure 5d.
[44] Several previous studies of basalt cores or logs have
used a cross plot of either apparent or intrinsic formation
factor versus porosity to estimate a and m. For example,
Pezard [1990] found from cores at Hole 504B that F = 10
f1. For core and log data from Sites 768 and 770,
estimates of a range from 2.2 to 6.5, and estimates of m
range from 1.2 to 1.6 [Jarrard and Schaar, 1991]. For the
very old crust of Hole 418A, Broglia and Moos [1988]
found from logs that F = 11.5 f1.85 in the uppermost
altered interval, and F = 29.5 f1.16 in the lower portion of
that site. Similarly, the three major lithostratigraphic subdivisions (alkalic basalts, hydrothermal zone, and tholeiitic
basalts) of Hole 801C have subtly different relations
between apparent formation factor and porosity [Jarrard
et al., 1995]. For the tholeiites, the highest-porosity rocks
indicate a linear trend with a of 4 – 10, but low-porosity
zones exhibit little apparent correlation of formation factor
to porosity, probably because of porosity errors.
[45] Figures 6c and 6d compare Hole 801C apparent
formation factors to porosity for both core plug and log
data. Compared to the close relationship between these
parameters shown by core data, log data have high scatter.
We attribute this scatter to porosity errors, as was the case
for Leg 144 logs, due to inadequate contact of the density
tool pad with the borehole wall. If equilibrium temperatures
had been assumed for the calculation of apparent formation
factors from resistivities, rather than assuming some formation cooling by drilling circulation, the pattern in Figure
6d would be offset slightly upward; this uncertainty is minor
in comparison to the range of formation factors and their
scatter. For the core data, linear regression indicates that F =
2.6 f2.1, within the range of relationships observed at other
basalt sites. The log-based pattern, though high in scatter,
appears to be systematically offset to the left of the corebased pattern (lower porosities for a given formation factor).
[46] Figures 6e and 6f compare apparent formation factors to velocity, for both core plug and log data. Because
both depend on porosity, a positive correlation is expected.
Both core plug and log data do show a good correlation
between these parameters. Comparison of the two trends
indicates generally lower formation factors, for a given
velocity, for log data than for core data. Velocity difference
between in situ and laboratory pressures is unlikely to be
responsible for this offset between trends, because rebound
here is small [Wallick et al., 1992] and reduces both velocity
and formation factor.
[47] The cross plots of Figure 6 indicate that apparent
formation factor is the geophysical property that differs most
between core and log scales. This conclusion is confirmed
by overlays of core and log measurements for velocity,
density, and formation factor versus depth (Figure 2).
Apparent formation factors are lower for logs than for core
plugs. Either subvertical cracks (which are low resistivity
but not seen by velocity logs) or crack filling by clays
(which are conductive yet increase velocity) may explain
this difference.
[48] The relationships among log-based velocity, formation factor, and density seen in Figure 6 differ substantially
in data scatter. Velocity is much better correlated with
formation factor than either of these parameters is correlated
with density. Comparison of depth plots of the three logs
(Figure 2) confirms this conclusion. We attribute the density
discrepancies to artifacts within the density log, caused by
poor pad contact. Although the Hole 801C density log is
one of the best quality density logs obtained by either DSDP
or ODP from oceanic crust, its reliability is marginal.
Fortunately, velocity and resistivity logs are just as sensitive
to porosity variations as is the density log, and they have
much higher signal-to-noise ratios than does density logging
in basalts.
[49] Velocity/formation factor relationships for oceanic
basalts have not been analyzed previously, but some porosity/formation factor log analyses have used sonic-based
porosities because density-based porosities were poor or
unavailable [Jarrard and Broglia, 1991; de Larouzière et
al., 1996]. At log scale, the velocity/formation factor
relation may depend on alteration intensity [de Larouzière
et al., 1996]. For Hole 801C core plugs, however, the least
and most altered samples lie along the same velocity/
formation factor trend.
3.6. Potassium Enrichment
[50] Is the plug-based correlation between alteration and
porosity (Figure 5c) also evident in the continuous records
available from in situ logs? We investigated this question by
comparing log measurements of velocity and potassium
(Figure 7).
[51] The potassium content of altered basalts is often an
order of magnitude higher than that of fresh basalts, due to
potassium absorption from seawater during low-temperature
alteration [Hart, 1969]. Potassium-rich saponite and celadonite are commonly formed during low-temperature diagenesis [e.g., Alt and Honnorez, 1984]. Consequently,
Broglia and Moos [1988] found that both the potassium
log and total gamma ray log at Hole 418A were highly
correlated with qualitative ratings of core alteration, and
both indicate variation in content of potassium-rich alteration minerals at that site. A relationship between gamma
ray and either velocity, resistivity, or fracture intensity logs
has been observed at Holes 395A [Moos, 1990], 765D
[Shipboard Scientific Party, 1990a], and 896A [de Larou-
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
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5 - 13
imply a high proportion of potassium-bearing alteration
minerals; such zones are less likely to be recovered by coring.
This strong log-based correlation between porosity and
potassium-bearing alteration minerals is consistent with a
core-based correlation for the upper tholeiites [Busch et al.,
1992] and our core-based correlations between porosity and
hydrous minerals. Although the log-based pattern could be
generated by fluid flow and associated precipitation either in
cracks or at an intergranular scale, the core-based patterns
require intergranular flow and alteration.
Figure 7. Comparison of velocity and potassium logs for
Hole 801C tholeiites. The strong inverse correlation
between these logs indicates that precipitation of potassium-bearing clays and celadonite occurs mainly in the most
porous (lowest-velocity) zones. See color version of this
figure at back of this issue.
zière et al., 1996], but Sites 768 and 770 exhibit only a
subtle association between alteration indicators and potassium [Jarrard and Broglia, 1991].
[52] For the alkali basalts and upper tholeiites drilled at
Site 801 during Leg 129, a cross plot of K2O versus H2O+ is
a moderately successful alteration indicator [Alt et al.,
1992]; primary compositional variations among the alkalic
basalts and potassium-free alteration minerals degrade the
relationship. Jarrard et al. [1995] concluded that Th/K ratio
was superior to total potassium in detecting alterationinduced potassium enrichment among the upper tholeiites;
it too was nondiagnostic among the alkalic basalts. Comparison of the Th/K log to the porosity and velocity logs
suggested that greater alteration is correlated with both
greater porosity and lower velocity. Leg 185 tholeiite coring
demonstrates a strong correlation between patent alteration
in the core, K2O concentration based on atomic absorption
and XRF and high gamma ray counts measured both in
cores on the multisensor track and in situ with downhole
logs [Shipboard Scientific Party, 2000a]. For example, a
zone at 530 mbsf with K2O >2 wt % is celadonite-rich,
and a K2O minimum at 720 mbsf has abundant fresh glass
[Shipboard Scientific Party, 2000a]. Average K2O from
downhole logs is slightly higher than that from cores
[Shipboard Scientific Party, 2000a], probably because of
higher core recovery in fresh, massive units than in interpillow material and highly altered zones.
[53] Figure 7, which overlays potassium and velocity logs
versus depth for the Hole 801C tholeiites, demonstrates that
potassium variations are closely linked to porosity changes.
We use velocity as a porosity indicator, rather than density,
because of density log errors discussed above. The lowest
observed potassium contents, 0.06 – 0.1 wt%, are comparable to values typical of unaltered MORB tholeiites [Spivack
and Staudigel, 1994; Hart et al., 1999; Shipboard Scientific
Party, 2000a], indicating that the most massive, low-porosity
units are minimally altered. Potassium maxima of 1%
3.7. Synthesis
[54] Hole 801C tholeiites demonstrate the relationships
among hydrologic properties, geophysical properties, and
crustal alteration (Figure 8a). At an intergranular scale, high
porosity generates permeability that promotes fluid flux and
resulting alteration (Figure 8a). Alteration, in turn, may
increase intergranular porosity and permeability (Figure 8a).
Greater alteration, as indicated by H2O+ [Jarrard et al., 1995]
or LAS, is associated with both greater porosity (Figure 5c)
and permeability (Figure 4b). Alteration also reduces matrix
densities and matrix velocities by up to 10%. The relationship
between velocity and density is best fit by a model in which
higher porosities are associated with alteration-induced
reductions in matrix velocity and matrix density (Figure 6a).
[55] Whereas plugs reveal intergranular patterns, coring
misses cracks and crack filling that potentially control macroporosity and large-scale velocity [Hyndman and Drury,
1976]. Log analysis, in contrast, allows us to determine
whether the laboratory sample-scale observations also hold
at the interflow scale. Our comparison of Hole 801C log
velocity to density shows a pattern (Figure 6b) that agrees
with core data (Figure 5a). In contrast, the relationship of log
formation factor to velocity is offset from that seen at the
intergranular scale (Figures 6e and 6f), possibly because of
conductive clays in cracks. The potassium log confirms that
greatest alteration is correlated with highest porosities and
lowest velocities (Figure 7). The net effect is that the most
altered zones remain lower in velocity than the lowerporosity, lower-permeability, unaltered zones. An extreme
example is at 711 mbsf (Figures 2 and 7): this probable crack
still has the highest permeability, lowest velocity, and lowest
formation factor within the tholeiites, yet it has the most
hydrous minerals (based on the potassium log).
[56] These observations at Hole 801C lead to the following
tentative predicted effects of age-dependent crustal alteration
on evolution of geophysical properties: (1) ongoing alteration causes significant decreases in intergranular (core plug)
matrix density, matrix velocity, bulk density, and velocity,
possibly accompanied by subtle increase in intergranular
porosity; (2) alteration reduces macroporosity; (3) the alteration impact on log-scale velocity involves the competing
effects of fracture filling and intergranular velocity reduction;
(4) alteration increases potassium, at both core and log scales;
and (5) these patterns will be more evident in pillows than in
flows because of permeability differences.
4. Temporal Variations in Physical Properties of
Upper Oceanic Crust
[57] The evolution of hydrothermal circulation is well
established qualitatively. Hydrothermal circulation is most
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5 - 14
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
Figure 8. Flowcharts of the interrelations among hydrologic and geophysical parameters during the aging and
alteration of oceanic crust. (a) Observed pattern at Hole
801C, based on core plug measurements. (b) Generally
accepted ‘‘standard model.’’ (c) Expected relationships at
the interflow scale, based on the standard model; some links
have been observationally verified, whereas others are only
assumed. Note that alteration increases large-scale velocities
by reducing macroporosity, whereas it decreases intergranular velocity. Modified from Jarrard et al. [1995].
vigorous and extends deepest into the oceanic crust at the
ridge axis, where thermally induced buoyancy forces are
highest [Humphris, 1995; Fisher, 1998]. If early sedimentation is substantial, hydrothermal circulation is isolated
from seawater and higher in temperature [Fehn and Cathles,
1986; Davis et al., 1992; Humphris, 1995]. Hydrothermal
circulation continues within ridge flanks for at least tens of
millions of years [e.g., Anderson et al., 1977; Fisher et al.,
1990]. Off-axis waning of hydrothermal circulation is
caused by decrease in crustal heat, accumulation of a
relatively low-permeability sediment cover, and decrease
in crustal permeability due to crustal alteration [Anderson et
al., 1977; Stein and Stein, 1994]. Of these three mechanisms
by which hydrothermal circulation is reduced with time and
distance from the ridge, crustal alteration has received the
most attention, perhaps because only it can account for the
age dependence of upper crustal velocities [Schreiber and
Fox, 1976, 1977]. Honnorez [1981], Thompson [1991], and
Alt [1995] provided comprehensive reviews of low-temperature crustal alteration processes. Alteration minerals eventually fill cracks and interpillow voids, thereby reducing
porosity and permeability and increasing velocity (Figure
8b). This model suggests that hydrothermal circulation is
ultimately self-limiting. It is, however, opposite to the
pattern observed at Hole 801C (Figure 8a) in a fundamental
respect: greater alteration at Hole 801C is associated with
high porosity and low velocity. This contradiction may be
only apparent: the highest-porosity zones probably experience the most alteration and associated porosity reduction,
yet they remain more porous than fresher, more massive
intervals.
[58] Of the many associations that are expected to result
from crustal aging, some are only assumed rather than
directly observed (Figure 8c), and others have been
observed only in young crust. The correlation between
greater alteration and greater velocity, assumed to be present
at the multiflow scale (Figure 8c), has only been demonstrated at one site, Hole 418A [Broglia and Moos, 1988].
Alteration, though extremely heterogeneous, certainly does
generally increase with crustal age very near the ridge crest,
and abundance of crack-filling carbonate veins is higher in
old crust than in 8 Ma crust [Alt and Teagle, 1999]. Hole
504B provides direct evidence for alteration-induced permeability reduction [Anderson and Zoback, 1982; Anderson
et al., 1985; Becker, 1989]. Bulk permeability decreases
systematically with crustal age, for the first 8 Myr [Fisher
and Becker, 2000], but packer experiments at Hole 801C
found some of the highest crustal permeability ever
observed [Larson et al., 1993]. To evaluate both the generality of Hole 801C observations and the inferred patterns
of crustal aging, it is necessary to undertake a global
examination of temporal variations in physical properties
of upper oceanic crust.
4.1. Previous Evidence for Timing of Waning of
Hydrothermal Circulation
[59] The temporal evolution of hydrothermal circulation
in oceanic crust is an outstanding problem. Different geophysical techniques suggest that ridge flank hydrothermal
circulation terminates at 5– 10 Ma, >10 Ma, 10– 40 Ma,
65 Ma, 100 Ma, or beyond the Cretaceous.
[60] Global data syntheses of upper crustal seismic velocities [Grevemeyer and Weigel, 1996; Carlson, 1998] indicate that velocities increase approximately linearly with the
logarithm of time from 0.1 Ma to 5– 10 Ma, but changes
after that are subtle or absent. Grevemeyer et al. [1999]
demonstrate this systematic increase within a single 0 – 8 Ma
heat flow and seismic velocity transect on the west flank of
the East Pacific Rise.
[61] Bulk permeabilities decrease exponentially from 1 to
8 Ma and may continue to decrease, though subsequent
changes are undetermined [Fisher and Becker, 2000]. Most
of these permeability data are from sedimented ridges,
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
which may have alteration histories that are not representative of open ocean crust.
[62] Dating of alteration minerals within upper oceanic
crust suggests that alteration terminates by 10– 40 Ma
[Peterson et al., 1986; Gallahan and Duncan, 1994; Hart et
al., 1994]. Most of these dates are <15 Myr younger than
the age of surrounding crust, but some indicate that formation of alteration minerals continues for 40 Myr [Gallahan and Duncan, 1994].
[63] The global heat flow synthesis of Stein and Stein
[1994] shows that the heat flow deficit (theoretical minus
observed) gradually decreases with age, disappearing at
about 65 Ma. The deficit indicates advection via open
circulation between seawater and the crust, with narrow
unsampled zones of upwelling heat loss. Beyond 65 Ma,
heat flow is dominantly conductive. The sealing age
depends primarily on crustal age and only secondarily on
sediment thickness, implying that it is caused mainly by
decreased crustal permeability and/or reduced temperature
contrast [Stein and Stein, 1994; Stein et al., 1995]. Independent evidence of hydrothermal circulation in 15– 70 Ma
crust comes from geochemical gradients within pore fluids
of equatorial Pacific sediments [Baker et al., 1991].
[64] Velocities and densities of basalt core plugs decrease
as crust ages, until 100 Ma [Johnson and Semyan, 1994].
This analysis of global DSDP (and minor ODP) data from
the top 50 m of oceanic crust confirmed the much earlier
Christensen and Salisbury [1973] velocity results for the
uppermost 1 – 2 m of basalts.
[65] Hydrothermal circulation may terminate beyond the
Cretaceous, if at all, based on heterogeneity within some
regional heat flow surveys on Cretaceous crust [e.g.,
Embley et al., 1983; Noel and Hounslow, 1988]. Average
heat flow in these regions is conductive, but variations
within each region suggest hydrothermal circulation. If
global heat flow data are binned by crustal age, percentage
standard deviation decreases systematically with age, even
within the Mesozoic [Stein and Stein, 1994]. This observation indicates that waning hydrothermal circulation is
still present in some old crust [Stein and Stein, 1994].
Although heat flow data cannot unambiguously determine
whether this hydrothermal circulation in old crust is closed
cell (no interchange with seawater) or slow open cell
[Fisher, 1998], the two differ crucially in geochemical
mass balance. For closed cell fluid flow, precipitation
equals solution, and crustal geophysical properties may
be unchanged by ongoing alteration.
[66] These conflicting indications of when off-axis hydrothermal circulation wanes have been partially reconciled.
Apparently, any crustal alteration that may occur beyond
5 – 10 Ma is so subtle that some geophysical methods may
be incapable of resolving it. Fisher and Becker [2000]
suggest that the rapid initial increase in velocity and
decrease in permeability are not necessarily incompatible
with heat flow evidence for persistence of open cell convection to 65 Ma. They hypothesize that hydrothermal
circulation continues to 65 Ma, in localized channels within
the crust that are too small a percentage of total upper crust
for significant effects on average crustal velocity. If the heat
flow evidence for continued fluid flow could be coupled
with a demonstration of the magnitude of ongoing crustal
alteration, estimates of global geochemical fluxes into the
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5 - 15
oceans [Kadko et al., 1995; Elderfield and Schultz, 1996]
and subduction zones could be refined. Sections 4.2 –4.5
synthesize core, log, and seismic refraction evidence of
whether physical properties of upper oceanic crust continue
to evolve beyond 5 Ma.
4.2. Crustal Velocity Versus Age
[67] Houtz and Ewing [1976] first demonstrated that
sonobuoy refraction velocities for layer 2A increase systematically with increasing crustal age. This observation was
attributed to lowering of uppermost crustal macroporosity
by hydrothermal alteration products (Figure 8b). Subsequent refraction experiments using improved techniques
demonstrate that seismic velocities of the upper oceanic
crust increase rapidly for the first 5 –8 Myr, from an average
of 2.3 km/s at the spreading center to 4.3– 4.4 km/s at 5 – 10
Ma [Grevemeyer and Weigel, 1996; Carlson, 1998]. Velocity modeling suggests that the large near-ridge velocity
increases may be accomplished by changing the shape of
pore spaces via secondary mineralization with relatively
small overall porosity reduction [Wilkens et al., 1991; Shaw,
1994; Moos and Marion, 1994]. Beyond 10 Ma, no significant change in upper crustal velocity is evident.
[68] The uppermost layer of the seismic modeling studies,
which is 200– 500 m thick, is the portion of oceanic crust
that exhibits the greatest age-dependent change [White,
1984]. Thus the depth interval of the DSDP and ODP sites
should provide a sensitive barometer of crustal aging for all
crustal ages. Figure 9 plots average velocity versus logarithm of crustal age, based on averaging each of the 13
velocity logs in Figure 3. The velocity averages shown are
actually based on averaging transit times (the inverse of
velocity) rather than velocity because the seismic method
similarly averages transit times rather than velocities. The
95% confidence limits shown are optimistic, because the
assumption that each log value is independent is invalid in
depth series. Intersite variation is high because of local
variations in degree of alteration, initial porosity, and
volcanic style.
[69] Average velocity for the flows ranges from 4.6 to 5.7
km/s, with no age dependence of velocity evident (Figure 9).
We attribute this lack of age dependence to initial porosities
and permeabilities too low for significant fluid flow and
associated macroporosity reduction. Average velocity for
the pillows appears to be 4.25 –4.7 km/s initially, increasing
with age to 4.8 – 5.0 km/s. Scatter at intermediate ages
encompasses nearly the full range of observed velocities,
precluding detection of any changes after 30 Ma. Overall
average velocity, like pillow velocity, is initially 4.2 –4.7
km/s and increases with age to 5.0– 5.1 km/s. This change
appears to continue beyond 30 Ma, but the evidence is not
compelling. Because pillows are generally much more
abundant than flows, the patterns for pillows and total
basalts are mostly similar. Hole 504B (the youngest site),
however, has many high-velocity flows and so is anomalously fast in total average velocity.
[70] The age dependence of total average velocity shown
here (Figure 9) extends the already well-established age
dependence of uppermost crustal velocities for <5– 10 Myr
based on seismic studies [Houtz and Ewing, 1976; White,
1984; Grevemeyer and Weigel, 1996, 1997; Carlson, 1998;
Grevemeyer et al., 1999]. Individual seismic studies, though
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5 - 16
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
plug velocities are higher than log velocities (Figure 3)
because interpillow porosity and cracks are seen only by the
logs. Velocity logs indicate slightly higher velocities, by
0.2 – 0.3 km/s, than the seismic results, as described
above. This difference is likely to be open fracture porosity
in large, near-vertical fractures and normal faults, seen by
the seismic experiments but not by vertical logs.
[72] Crustal age and velocity are positively correlated at
the multiflow scale, but inversely correlated at the plug scale.
Christensen and Salisbury [1972, 1973] first observed that
core samples from the uppermost 2 m of oceanic crust
typically exhibit a velocity decrease associated with
increasing crustal age. Low-temperature alteration breaks
down glass, olivine, and feldspar, and replaces them with
clay minerals that are lower in matrix velocity and density.
Potassium, sodium, and water are absorbed from seawater;
calcium, silicon, and iron are given off, and the net
geochemical flux is out of the basalt [e.g., Hart, 1973],
thereby increasing porosity at an intergranular scale. The
plug-based relation of greater alteration inducing lower
velocity apparently is competing with alteration-induced
crack filling and large-scale velocity increase. Consequently, core and log velocities converge as crust ages
(Figure 3). This age-dependent convergence is examined in
section 4.3.
Figure 9. Log-based velocity as a function of crustal age,
for the 13 sites of Table 2, compared to the seismic velocity
synthesis of Carlson [1998]. The moderate log-based trend
toward velocity increase with increasing age (top) for total
basement rocks results from (middle) the age dependence
for pillows. (bottom) In contrast, flows exhibit no agedependent change.
subject to dip and other analytical uncertainties, are both
more abundant and average larger basement volumes than
the individual sites of this study. Nevertheless, the logs of
this study augment the seismic results in two critical ways.
First, this study is able to link velocity change to volcanic
style and macroporosity reduction. Second, the logs appear
to indicate systematic velocity increase for the time period
6 – 167 Ma. Linear regression of log results for this time
period indicates that average upper crustal velocity
increases by 0.5 km/s, from 4.5 to 5.0 km/s (Figure 9).
Regression of Carlson’s [1998] compiled seismic results for
this period suggests slightly lower velocities than the log
results but a similar velocity increase versus age. However,
scatter is much higher for the seismic results than for log data
and the apparent increase in seismic velocity is not significant at the 95% confidence level. Consequently, Carlson
[1998] concluded that rapid velocity change is limited to the
first 5 –10 Myr and that no significant increase is detected
beyond that age.
[71] Upper crustal velocities appear to be scale-dependent, decreasing from core plug to log to seismic scale. Core
4.3. Macroporosity Versus Age
[73] Plugs reveal intergranular patterns, but they tend to
miss cracks and crack-filling that potentially are major
controls on large-scale velocity [Hyndman and Drury,
1976; Anderson and Zoback, 1983]. Estimation of crack
porosity and its effect on large-scale physical properties
using incomplete core recovery is difficult [Johnson,
1979b]. This crack filling may be more important to
large-scale crustal velocity than the opposing intergranular
effects of alteration. If much of basalt porosity is in the form
of large-scale voids and fractures that are not sampled by
10-cm3 core plugs, then comparison of core and log
measurements has the potential of detecting this macroporosity. We use velocity logs rather than density logs for
macroporosity detection because even the best density logs
have abundant artifacts due to tool sensitivity to borehole
irregularities (e.g., compare the Hole 801C density, velocity,
and formation factor logs in Figure 2). Comparisons of core
and log velocities are degraded, however, by low core
recovery and nonrepresentative core recovery. Core recovery in basalts varies from 10% to 80%, and the coring
process can bias basalt core recovery toward fresher, less
altered basalts. This reduced recovery leads to 0.5– 9.0 m of
depth uncertainty in attempts to match core and log measurements for the same depth.
[74] Many studies have compared basalt core and log
velocities at individual sites. Such comparisons at several
younger sites [Kirkpatrick, 1978; Moos et al., 1990; Jarrard
and Broglia, 1991] concluded that macroporosity missed by
the plugs caused lower velocity and density values for logs
than for plugs. For old sites, in contrast, the standard model
predicts that macroporosity is reduced or absent and therefore that plug and log velocities will be similar. This
prediction is confirmed at Holes 765D [Shipboard Scientific
Party, 1990a] and 801C (Figure 2), refuted at Hole 417D
[Salisbury et al., 1979], and dominantly refuted at Hole
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
418A [Moos et al., 1990]. Jacobson [1992] compared plug
velocities to nearby seismic reflection and refraction velocities and found that the difference between the two
decreased with increasing age. He attributed this convergence to filling of cracks by alteration minerals.
[75] Before comparing core plug and log velocities for
macroporosity determination, we must evaluate pressure
effects on velocity. Log velocities are measured in situ,
but we compare them to core velocity measurements made
at atmospheric pressure and reported by the shipboard
scientific parties. Rebound, the core expansion that accompanies change from in situ to laboratory pressure, can
reduce plug velocities [e.g., Nur, 1971; Bourbié et al.,
1987]. Rebound is generally minor in oceanic extrusive
basalts, because burial depths are much too small to induce
microcracks. The extensive suite of core velocity measurements for Hole 504B [Wilkens et al., 1983; Christensen and
Salisbury, 1985; Christensen et al., 1989; Iturrino et al.,
1995; Salisbury et al., 1996] shows that the average
percentage difference between atmospheric and in situ
velocities for the upper 600 m of basement is only 0.3%
based on differential pressures, 0.6% based on lithostatic
pressures, and 2.0% based on confining pressures. In contrast, the two diverge at greater depths: at 1370 mbsf,
600 m into the dikes, the unmistakable signature of
microcracks appears and atmospheric pressure measurements are no longer a useful indicator of in situ velocities.
Consequently, we assume that atmospheric pressure velocities can be used for macroporosity calculations within this
and other extrusive sections, whereas high-pressure measurements would be required for dikes and gabbros.
[76] Another factor that can cause significant differences
between core and log velocities is dispersion [Spencer,
1981; Bourbié et al., 1987; Goldberg and Zinszner, 1988],
which generates higher velocities for ultrasonic frequency
measurements on cores than for much lower-frequency log
measurements. The excellent agreement of log and highpressure core velocities within the dikes at Hole 504B
[Salisbury et al., 1996], where macroporosity is near zero,
indicates that dispersion introduces little bias into basalt
core/log velocity comparisons. Furthermore, dispersion
effects are unlikely to be age-dependent.
[77] Figure 10 compares core plug and log velocities for
the 13 sites of this study. Because a systematic difference
between the two is thought to result primarily from the
influence of macroporosity on logs but not on cores, we
have expressed the difference between each core/log velocity pair as an apparent porosity difference L C =
100*(tL tC)/(tf tma), where L is percentage
porosity calculated from log velocity, C is percentage
porosity from core plug velocity, tL is measured log
transit time (inverse of velocity), tC is measured core
transit time, tf is fluid transit time (0.62 s/km), and tma
is matrix transit time. Matrix transit time is assumed to be
0.147 s/km, the value for fresh basalt. Alteration actually
reduces tma by up to 10% [e.g., Jarrard et al., 1995], but
the resulting macroporosity error of <3% of calculated
values is minor in comparison with the order of magnitude
macroporosity variation among sites. Apparent macroporosity (L C) is plotted as a function of age in
Figure 10, for pillows, for flows, and for all basement rocks
combined. For Holes 395A, 418A, and 504B, the robust-
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5 - 17
Figure 10. Apparent macroporosity (defined as the
difference between plug-scale and log-scale porosities
calculated from velocity measurements) as a function of
crustal age for the 13 sites of Table 2. The pattern of
macroporosity reduction with increasing age (top) for all
basement rocks results from (middle) the age dependence
for pillows. (bottom) In contrast, flow macroporosity is
significantly lower and exhibits no systematic change with
age.
ness of macroporosity values could be evaluated by comparing results of Schlumberger and multichannel sonic logs:
calculated macroporosities differed by +2.0%, 0.6%, and
+0.1%, respectively.
[78] Flows indicate little or no systematic change of
macroporosity with increasing crustal age. Macroporosity
is so low (usually <4%) for all crustal ages that only half of
the site averages are significantly greater than zero at the
95% confidence level. Core descriptions routinely refer to
flows as massive, with porosity generally confined to
abundant cracks immediately adjacent to the flow boundaries. Undoubtedly this crack porosity does undergo some
filling during alteration, but this surface change probably
has minimal effect on average porosity of a flow.
[79] Pillow macroporosities are substantially higher than
flow porosities, and they exhibit a decrease with age (R =
0.89) that is significant at the 99% confidence level,
dropping from an initial value of 10% to 0 – 4% beyond
30 Ma (Figure 10). These data do not demonstrate con-
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JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
vincingly that porosity reduction continues beyond 30 Ma,
though the two oldest sites have the lowest pillow macroporosities. The pattern of total macroporosities is very
similar to that from pillows, because pillows are the
dominant component at most sites. Total macroporosities
also exhibit a decrease with increasing age (R = 0.85)
that is significant at the 99% confidence level, and this
decrease appears to persist rather than wane at 30 Ma.
Initial macroporosities of 8 – 10% eventually are reduced to
near zero. This pattern is a strong confirmation of the
standard model, which predicts age-dependent macroporosity reduction.
[80] For the 13 sites of Table 2, core plug velocities
exhibit an inverse correlation with logarithm of age that is
significant at the 95% confidence level, both for total basalts
(R = 0.57) and for pillows (R = 0.68). Core plug
velocities of flows are not significantly correlated with
logarithm of age (R = 0.26). Temporal increase in overall
large-scale velocity due to fracture and void filling occurs
despite alteration-induced decrease in core plug velocity.
We conclude that the gradual convergence of core and log
velocities (Figure 3) is not simply a macroporosity decrease
(Figure 10) but a combination of fracture and void filling
(increased log velocity) and decreased core plug velocity.
Consequently, the age dependence of apparent macroporosity (Figure 10) is stronger than that for either log velocity
(Figure 9) or core plug velocity, despite depth and volume
mismatches inherent in comparing log and core plug velocities to calculate macroporosity.
4.4. Potassium Versus Age
[81] Low-temperature crustal alteration processes include
oxidation, hydration, and alkali fixation. The latter two
generate clay minerals, the most abundant alteration products [e.g., Donnelly et al., 1979; Gillis and Robinson, 1988;
Alt and Honnorez, 1984; Alt et al., 1986]. Consequently,
altered basalts are generally higher in potassium than
unaltered basalts [Hart, 1969; Hart et al., 1974]. The extent
and timing of alteration-induced potassium enrichment of
oceanic crust has implications for the potassium budgets of
the ocean and subduction zones. Available data, however,
have been generally sparse and nonrepresentative. Most
major element analyses of DSDP and ODP basalts select
the least altered samples, because the geochemical objective
is original composition rather than alteration-induced
changes. Basalt core recovery is biased toward less altered
materials, with underrepresentation of hyaloclastites and
interpillow alteration minerals.
[82] Spectral gamma ray logs, which provide continuous
records of potassium concentration, offer additional perspectives on alteration-induced potassium enrichment. The
potassium log at Hole 418A is highly correlated with
qualitative ratings of core alteration [Broglia and Moos,
1988], whereas potassium logs for Sites 768 and 770 are
weakly correlated with alteration indicators [Jarrard and
Broglia, 1991]. We interpreted the strong inverse correlation
between potassium and velocity logs within Hole 801C
tholeiites (Figure 7) as indicating production of high-K
clays and celadonite primarily in the most porous and
therefore most permeable zones, possibly augmented by
alteration-induced increase in intergranular porosity and
decrease in matrix velocity.
Figure 11. Relationship between core-based matrix density and log-based potassium for sites of Table 2. Hydration,
the dominant type of crustal alteration, lowers matrix
density and absorbs potassium from seawater. Solid dots
indicate individual sites.
[83] Only seven of the 13 logged basement sites in
normal oceanic crust (Table 2) have potassium logs. DSDP
logging used a total gamma ray tool, whereas ODP logging
has routinely employed spectral gamma ray logging; fortunately, several DSDP basement sites have been revisited
by ODP and logged with newer tools such as spectral
gamma ray. Among these seven sites, average log-based
potassium content is strongly correlated with average corebased matrix density (R = 0.86; significant at 95%
confidence level) (Figure 11). This correlation demonstrates
that alteration of upper oceanic crust entails potassium
enrichment. In contrast, the high-temperature alteration of
deeper oceanic crust that occurs at the ridge crest depletes
potassium [Elderfield and Schultz, 1996; Kadko et al.,
1995].
[84] The few normal crustal sites with potassium logs
may indicate that old (>100 Ma) sites are higher in
potassium than young (6 –7 Ma) sites (Figure 12), but the
correlation (R = 0.47) is not statistically significant, and
potassium-depleted magmas affect the values for two of the
youngest sites (504 and 896) [Alt et al., 1996]. Potassium
content of pillows is usually slightly higher than that of
flows from the same site (Figure 12 and Table 2), as
expected because of the greater permeability and alteration
of pillows. The highest-potassium site is highly altered Site
770, which is backarc-trapped crust of intermediate age.
The correlation without the extreme value at Site 770 is
much greater: R = 0.64 for all units, R = 0.65 for flows, and
R = 0.58 for pillows.
4.5. Core Alteration, Bulk Density, Porosity, and
Matrix Density Versus Age
[85] Basalt core descriptions routinely include subjective
rankings of alteration intensity. Criteria differ from leg to
leg, and no systematic pattern of increasing alteration rank
with increasing age is evident. Progressive CO2 enrichment
of oceanic crust has been quantified by analysis of five
DSDP and ODP sites [Staudigel et al., 1996; Alt and Teagle,
1999]. Alt and Teagle [1999] used these results to demon-
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
Figure 12. Potassium as a function of crustal age, for the
seven sites of Table 2 that have spectral gamma ray logging.
Age-dependent crustal alteration may cause progressive
potassium enrichment, but the pattern is not statistically
significant.
strate that carbonate precipitation in veins continues beyond
the 6 Ma age of Sites 504 and 896, resulting in enrichment
of total CO2 content of the upper extrusive section.
[86] Johnson and Semyan [1994] examined the possibility of age-dependent changes in velocity, porosity, and
density, based on a compilation of core plug measurements
from DSDP and ODP basalts. Far more drill holes into
basalt had velocity (65 holes) or density (50 holes) data than
porosity data (25 holes), and only two of these, including
Site 801, were ODP sites. Johnson and Semyan [1994]
found that average velocity and bulk density decreased
systematically from 0 to 100 Ma, but data older than 100
Myr were inexplicably highly scattered and often higher
than the trend for younger crust. Porosities exhibited a
possible increase with age, but the change was not statistically significant. Note that these patterns of age-dependent
change at the plug scale are the opposite of those seen at log
scale (e.g., Figures 9 and 10).
[87] Johnson and Semyan [1994] also compiled averages
of iron oxidation ratio and of H2O+ from major element
analyses. They found that oxidation occurs mainly within
the first 5 Myr and exhibits no detectable increase beyond
20 Ma. Zhou et al. [2001] measured oxidation state of
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5 - 19
titanomagnetites in 97 MORB samples and demonstrated
that small grains exhibit a strong linear pattern of increasing oxidation with logarithm of age. Their independent
data confirm the findings of Johnson and Semyan [1994]:
age-dependent oxidation is clearest for 0.1 –5 Ma and is
not resolved beyond 10– 30 Ma. The H2O+ compilation
of Johnson and Semyan [1994] confirms earlier indications [Hart, 1970, 1973; Donnelly et al., 1979; Muehlenbachs, 1979] that hydration increases with increasing age.
Unfortunately, the timing of this change is ill-determined:
H2O + correlates better with linear age (significant at 95%
confidence level) than with logarithm of age (not significant), whereas oxidation correlates better with logarithm
of age (significant at 99% confidence level) than with
linear age (not significant). Both velocity and density
decrease with increasing H2O+ (both significant at 99%
confidence level) [Johnson and Semyan, 1994], probably
because of the combination of porosity-dependent alteration and alteration-induced decrease in matrix density and
matrix velocity.
[88] Dating of alteration minerals, particularly celadonite,
commonly yields ages 8 – 16 Myr younger than the age of
the crust [Peterson et al., 1986; Gallahan and Duncan,
1994; Hart et al., 1994]. The comprehensive dating of
celadonites within the Troodos ophiolite indicates that half
of these alteration minerals formed more than 12 Myr after
the crust, and alteration continued for at least 40 Myr
[Gallahan and Duncan, 1994], particularly in the more
permeable units. However, discrepancies between Rb/Sr
and K/Ar ages and between core and outcrop ages suggest
that these durations could be overestimates [Booij et al.,
1995]. Furthermore, ophiolite alteration is nonrepresentative: it is generally more intense, with much higher fluid
flux, than normal crustal alteration [Alt and Teagle, 2000].
[89] Our compilation of temporal changes in core plug
physical properties draws on almost the same DSDP and
ODP databases as those used by Johnson and Semyan
[1994]; only five additional ODP basalt sites have obtained
significant penetration of normal oceanic crust and associated index property measurements. Both compilations are
confined to normal, open ocean oceanic crust. Johnson and
Semyan [1994] also excluded sites with anomalous depths,
whereas we exclude data from off-axis volcanism and
sedimented ridges. They included sites with any basement
penetration and data from only the top 50 m of basement.
We only include sites with >40 m of basement penetration,
in order to obtain a more representative crustal sample, and
we average results from the top 300 m of basement. Their
minimum number of measurements per site was two, but
ours is five, again to obtain a more representative sample.
They considered multiple holes at a site as independent
samples and calculated an average for each, whereas we
combine data from adjacent holes, except in rare cases
where the holes were drilled into different hydrothermal
environments (e.g., topographic high versus saddle). Both
compilations are subject to similar biases: preferential core
recovery and changing measurement techniques. For example, incomplete drying biases matrix densities and porosities
downward [Jarrard and Broglia, 1991; Jarrard et al.,
1995], whereas the routine technique of drying samples
by heating to 100– 110C biases them upward by driving
off interlayer water from clays.
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JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
Figure 13. Bulk density, porosity, and matrix density as a
function of crustal age based on core index property
measurements for the sites of Table 3. The weak pattern of
gradual density decrease results from systematic decrease in
matrix density, not from porosity increase. Arrows associated with the youngest site indicate that its actual age is
less than shown; its true age of 0.0 Ma cannot be plotted on
a logarithmic scale nor be used in the regressions.
[90] Unlike Johnson and Semyan [1994], we have not
compiled core plug velocities, except for logged sites as
described earlier. Our major focus is on matrix density, not
compiled by them, because our analyses of Site 801 and
other basalt sites led us to hypothesize that matrix density
measurements may provide sensitive and abundant indicators of crustal hydration. Figure 13 shows variations of bulk
density, matrix density, and porosity as a function of age,
based on site averages of core plug analyses (Table 3).
[91] No systematic change in porosity is evident, probably because of the competing effects of porosity filling by
alteration minerals and porosity generation during alteration
of minerals such as olivine. This observation at the plug
scale does not imply lack of age-dependent decrease in
macroporosity.
[92] Bulk densities may decrease with increasing age (R =
0.22; not statistically significant) (Figure 13). The correlation is much higher (R = 0.41, significant at 95% confidence
level) if low-porosity, low-density Site 410 is excluded. The
density decrease appears to continue throughout the
sampled age range of 0 – 167 Ma. This observation is
consistent with the systematic decrease found by Johnson
and Semyan [1994] for 0 – 100 Ma, but their observations of
anomalously high scatter and high densities beyond 100 Ma
are not confirmed by these data. The oldest three points on
this plot are new data from ODP Leg 185: Hole 801C, Hole
1149B, and Hole 1149D. The pattern of bulk density
decrease without systematic porosity change implies that
the source of density decrease is matrix density.
[93] Matrix density exhibits a strong pattern (R = 0.57;
significant at 99% confidence level) of decrease with
increasing age (Figure 13). The simplest interpretation of
this pattern is that matrix density continues to decrease, at a
rate proportional to logarithm of age, throughout the period
0– 167 Ma. However, data scatter in the time period 10– 70
Ma is high, and the hypothesis of no change beyond about
40– 60 Ma cannot be rejected. Index property sampling at
Site 597, which has an anomalously high average matrix
density (Table 3), is described as concentrating on a massive
flow [Shipboard Scientific Party, 1985c]; excluding this
point raises the correlation coefficient to 0.70.
[94] The matrix density plot of Figure 13 fails to remove a
major source of variance: Intersite variations in basalt
porosity affect permeability and therefore degree of alteration. Massive flows are usually much less altered than
porous pillows, and several studies of individual sites have
found a correlation between increasing porosity and
decreasing matrix density [e.g., Christensen et al., 1980;
Carlson and Herrick, 1990; Jarrard and Broglia, 1991;
Busch et al., 1992; Jarrard et al., 1995]. Figure 14 plots
site-average matrix densities for porosities of <5%, 5 – 10%,
10– 15%, and >15%. Sites less than 10 Myr old exhibit little
or no evidence of porosity-dependent reduction in matrix
density, for these grouped data. The same may be true for
15– 42 Ma sites, but near absence of porosities greater than
10% limits resolution in this time bracket. Older sites, 80–
167 Ma, have matrix densities that clearly decrease with
increasing porosity; among these are Leg 185 Holes 801C,
1149B, and 1149D.
[95] To reduce the influence of porosity on the pattern of
matrix density change due to crustal aging, we confine the
averaging of matrix densities to samples with a porosity of
>5% (Figure 15). Most massive flows, which are too
impermeable to share in progressive alteration of oceanic
crust, are thereby excluded. The pattern is quite similar to
that for all porosities, but the dependence on logarithm of
age is stronger (R = 0.72). The hypothesis of negligible
crustal alteration beyond 10 Ma [Carlson, 1998; Grevemeyer et al., 1999] can be rejected at the 99% confidence
level (R = 0.73 among sites >10 Ma).
4.6. Local Heterogeneity
[96] Superimposed on the age-dependent trends in logscale velocity (Figure 9), macroporosity (Figure 10), potassium (Figure 12), density (Figure 13), and matrix density
(Figures 13 and 15) is substantial local heterogeneity. This
heterogeneity, evident as scatter on these plots, is particularly evident at intermediate ages. Alteration appears to
occur earlier at some sites than others, though eventually all
are altered (Figures 9 and 13). Most of our 15– 40 Ma points
come from Leg 82 drilling in an area of the North Atlantic
where both alteration and geophysical properties are clearly
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
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5 - 21
1990, 1994]. This pattern is evident for paired holes on
adjacent topographic highs and saddles, at 417A/417D/
418A and at 504B/896A [Donnelly et al., 1979; Shipboard
Scientific Party, 1993; Alt et al., 1996].
[97] Heterogeneity of crustal alteration intensity hampers
efforts to determine the pattern of waning hydrothermal
circulation. Lack of logged sites prior to 6 Ma is not a
serious disadvantage, because ample seismic experiments
demonstrate the rapid velocity increase between 0 and about
5– 10 Ma [Carlson, 1998; Grevemeyer et al., 1999]. Paucity
of logged sites with ages between 15 and 100 Ma is
particularly unfortunate: changes in velocity (Figure 9),
macroporosity (Figure 10), and potassium (Figure 12)
within this time interval are poorly determined, and only
macroporosity suggests that upper crust of this age is
systematically less altered than crust >100 Ma. Fortunately,
this time interval is better represented by index property
measurements. Matrix densities of high-porosity basalts
provide the most comprehensive demonstration that crustal
alteration continues, albeit at a decreasing rate, throughout
the period 0 – 167 Ma, corresponding to virtually the entire
range of extant oceanic crust.
Figure 14. Relationships between porosity and matrix
density based on core index property measurements for the
sites of Table 3. High porosity permits increased fluid flow
and resulting increased alteration, but this effect is only
evident at older sites. See color version of this figure at back
of this issue.
heterogeneous. ODP Leg 187, which obtained basement
penetration of 7– 67 m at 13 sites south of Australia, is not
included in our figures because no logging or physical
properties measurements were undertaken. Visual core
descriptions of basalt alteration for these sites, ranging from
14 to 28 Ma, show high intersite heterogeneity and no
correlation between alteration and crustal age [Shipboard
Scientific Party, 2001]. The complexity of both vertical and
horizontal variations in alteration intensity is best documented in ophiolites, particularly the Troodos ophiolite
[Gillis and Robinson, 1988, 1990]. Basement topography
often controls the geometry of hydrothermal circulation,
concentrating outflow at topographic highs [Fisher et al.,
4.7. Rejuvenation of Hydrothermal Circulation at
the Outer Rise?
[98] The geochemical signature of age-dependent crustal
alteration can provide an essential starting point for calculations of geochemical mass balances associated with subduction. Implicit to such calculations, however, is the
assumption that normal oceanic crust is not modified by
the subduction process itself. This assumption may be
invalid.
[99] The flexure of lithosphere prior to subduction generates an outer rise. Extension at the top of the lithosphere
causes normal faulting, horst-and-graben topography [Ludwig et al., 1966], and associated seismicity with tensional
focal mechanisms [Stauder, 1968; Isacks et al., 1968;
Christensen and Ruff, 1988]. These phenomena imply
development and opening of near-vertical cracks. Within
ocean basins, off-ridge open cell hydrothermal circulation
appears to be limited largely by vertical permeability;
horizontal permeability is still relatively high [Fisher and
Becker, 2000]. Normal faulting at the outer rise might
rejuvenate hydrothermal circulation by generating vertical
permeability, particularly when combined with the strong
Figure 15. Matrix density as a function of crustal age,
based on core index property measurements for the sites of
Table 3. Because alteration is generally most intense in
more porous basalts (Figure 14), site-average matrix density
is calculated for samples with porosities of >5%.
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JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
Figure 16. MCS line 39, an east-west line recorded during Conrad Cruise 2005, illustrating the
proximity of Site 1149 to the Izu-Bonin Trench and the numerous normal faults cutting the sediment and
upper crust of the Pacific plate landward of the outer rise.
topographic gradient from the crest of the outer rise to the
trench floor. Strain cycling induced by cyclic seismic
flexure may also contribute to fluid flow [Silver et al.,
2000]. The very low heat flows measured immediately
seaward of the Costa Rica and Peru trenches may be
attributable to this mechanism [Langseth and Silver, 1996;
Silver et al., 2000].
[100] In addition to deepening Hole 801C, ODP Leg 185
drilled a deep crustal site on the outer rise of the Izu-Bonin
Trench (Figure 1). Holes 1149B and 1149D, though widely
spaced for holes from the same site, both encountered
severely altered basalts. Indeed, the two lowest matrix
densities on Figure 15 are these 132 Ma holes. Geochemical
gradients in interstitial waters just above basement at Site
1149 demonstrate that fluids are currently circulating within
this crust and, by geochemical mass balance, that this crust
is still undergoing alteration [Shipboard Scientific Party,
2000b]. Historic seismicity on the Izu-Bonin outer rise is
tensional [Christensen and Ruff, 1988; Kao and Chen,
1996]. Figure 16 shows that the crust surrounding Site
1149 is cut by abundant normal faults, as is typical for
outer rises.
[101] This correlation of crustal extension and faulting
with extreme alteration and active hydrothermal flow may
be coincidental. Alternatively, it may suggest that Site 1149
is representative of the final alteration state of subducting
normal oceanic crust. If so, estimates of geochemical mass
balance at subduction zones need to incorporate the effects
of hydrothermal rejuvenation within oceanic crust immediately seaward of the trench axis. Matrix densities for Site
1149 imply 25– 30% more hydration than is predicted by
the regression trend of Figure 15, but scatter from the trend
is high. Unfortunately, no other crustal sites have been
drilled on flexural outer rises. Additional heat flow surveys
might resolve this uncertainty, by testing the generality of
the pattern seen offshore Costa Rica by Langseth and Silver
[1996].
5. Conclusions
[102] ODP Hole 801C, which obtained the world’s oldest
section of in situ, normal oceanic crust, provides the
opportunity to examine relationships among hydrologic
properties (porosity, permeability, fluid flow), crustal alteration, and geophysical properties, at both core plug and
downhole log scales [Busch et al., 1992; Jarrard et al.,
1995]. Within these upper crustal basalts, higher porosity is
responsible for higher permeability and therefore higher
fluid flow rates. High fluid flux, in turn, fosters alteration,
particularly the hydration reactions that generate clays.
Consequently, porosity is well correlated with both LASbased hydration intensity (significant at 99% confidence
level; Figure 5c) and matrix density (significant at 99%
confidence level; Figure 5b). Because clay minerals are
20% lower in matrix density than the original basalt
constituents, hydration substantially reduces matrix density;
this effect is also demonstrated by the correlation between
matrix density and LAS-based hydration (significant at 99%
confidence level; Figure 5a).
[103] Geophysical properties such as formation factor are
impacted by this enhanced alteration in high-porosity zones,
resulting in a correlation between core plug formation factor
JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
and matrix density (significant at 99% confidence level;
Figure 5d). Porosity-dependent alteration is also seen at the
log scale: the excellent character match between potassium
and velocity logs (Figure 7) demonstrates that potassium
enrichment is proportional to porosity. Core plug results
indicate that hydration not only reduces density by reducing
matrix density, but also reduces velocity by reducing matrix
velocity. These combined effects are evident on velocity/
density cross plots, for both core plug (Figure 6a) and log
(Figure 6b) data.
[104] The similarity of core plug and log patterns at 801C
is significant, because the core plugs indicate intergranularscale properties, whereas the logs also respond to macroporosity. The log-based relationships of velocity to potassium (Figure 7) and of velocity to bulk density (Figure 6b)
could result from either intergranular-scale alteration or
hydration of cracks and interpillow voids. These two alteration scales cannot be isolated from the Hole 801C data
alone; a global examination of temporal changes in hydrologic, alteration, and geophysical properties is needed.
[105] To examine the evolution of physical properties for
normal oceanic crust, we used global data sets of DSDP and
ODP core physical properties and downhole logs (Figure 1
and Tables 2 and 3). These core and log data suggest that
crustal alteration continues, at a decreasing rate, throughout
the lifetime of oceanic crust. Crustal aging is accompanied
by increased hydration, increased log-scale velocity, and
decreased macroporosity.
[106] Matrix densities provide the strongest demonstration of systematic increase in alteration versus logarithm of
time (Figure 13). If variance associated with porositydependent alteration is reduced by confining the matrix
density analysis to porosities greater than 5% (Figure 15),
the ongoing evolution of matrix density is statistically
significant not only for the complete age interval but also
beyond 10 Ma (99% confidence level) and beyond 30 Ma
(95% confidence level).
[107] Large-scale velocity of the upper oceanic crust
increases with age. Seismic experiments demonstrate a
systematic increase of average velocity from 2.3 km/s at
the spreading center to 4.3– 4.4 km/s at 5 –10 Ma [Grevemeyer and Weigel, 1996; Carlson, 1998]. Beyond 5 – 10
Ma, in contrast, seismic velocities detect no significant age
dependence [Carlson, 1998], but DSDP and ODP log
velocities for pillows indicate that crustal velocity increase
continues (Figure 9; significant at 95% confidence level).
Both velocity increases are attributed to decrease in macroporosity; the former may be more rapid because of velocity
sensitivity to initial change in pore aspect ratio. This gradual
increase in large-scale velocity occurs despite systematic
velocity decrease at the intergranular scale, as indicated by
core plug velocities for 0 – 100 Ma [Christensen and Salisbury, 1973; Johnson and Semyan, 1994] and for the 13
logged basement sites. Precipitation in cracks and interpillow voids must be dominant over intergranular-scale alteration in controlling the evolution of large-scale crustal
velocity.
[108] Macroporosity can be estimated based on the difference between log and core plug velocity at the same depth.
Macroporosity decreases with increasing crustal age, with a
logarithm of age dependence that is evident in overall site
averages (significant at 99% confidence level; Figure 10)
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5 - 23
but even better demonstrated among the pillow basalts
(significant at 99% confidence level; Figure 10). This
systematic macroporosity decrease contrasts with the
absence of significant age-dependent porosity change
among core plugs, based on the compilations of Johnson
and Semyan [1994] and Figure 13.
[109] The time-dependent alteration of upper oceanic crust,
as well as associated changes in geophysical properties, can be
locally overwhelmed by two sources of variance: local topography, which controls hydrothermal circulation intensity,
and initial permeability, which is sensitive to porosity and
therefore also to volcanic style. Flows are generally much
more massive and impermeable than pillows, so they are
much less subject to alteration. Consequently, systematic agedependent changes in log-scale velocity, core plug velocity,
and macroporosity are much stronger for pillows (significant
at 95%, 95%, and 99% confidence level, respectively) than for
flows (none is significant). The observation of intense intergranular alteration of pillows and flow margins, despite
intergranular permeabilities that are more than 4 orders of
magnitude lower than the fracture and interpillow permeabilities measured by packers, requires that fluid flow and attendant alteration can proceed even at these exceedingly low
permeabilities. The lower limit for fluid flow and associated
alteration appears to be 1019 m2, based on the observation
that lower intergranular permeabilities are encountered
mostly in fresh, massive basalts with porosities of <5%
[Karato, 1983a, 1983b; Johnson, 1979a; Hamano, 1979].
Supporting evidence for this permeability threshold comes
from the lack of matrix density reduction for porosities of <5%
(Figure 13) and the lack of age dependence for both velocity
and macroporosity of massive flows (Figures 9 and 10).
[110] Examples of ongoing hydrothermal circulation in
very old crust, such as the present-day flow within basement
at Site 1149 [Shipboard Scientific Party, 2000b], microbial
alteration patterns within Hole 801C [Shipboard Scientific
Party, 2000a], and local heat flow anomalies in several
regions [e.g., Embley et al., 1983; Noel and Hounslow,
1988], do not appear to be confined to isolated instances.
Persistence of detectable hydrothermal circulation throughout the lifetime of oceanic crust is suggested by agedependent change in heat flow variance [Stein and Stein,
1994], velocity logs, macroporosity data, and matrix densities. This conclusion is not incompatible with previous
results suggesting lack of detectable change in seismic
velocity beyond 5 – 10 Ma [Carlson, 1998], preponderance
of alteration mineral ages < 40 Ma [e.g., Gallahan and
Duncan, 1994], and convergence of theoretical and
observed heat flows at 65 Ma [Stein and Stein, 1994].
The majority of velocity change and alteration is relatively
young, and termination of the heat flow deficit signifies a
reduction of advection to a level minor in comparison to
conduction, rather than an ending of hydrothermal flow
[Anderson and Skilbeck, 1981; Jacobson, 1992; Stein and
Stein, 1994]. On the basis of observed decreases in matrix
density that are proportional to the logarithm of age (Figures
13 and 15), approximately half of all intergranular-scale
crustal alteration occurs after the first 10– 15 Myr.
[111] Acknowledgments. This project was made possible by the data
acquisition efforts of shipboard scientific parties, operations managers, and
particularly the logging scientists of a dozen DSDP and ODP legs. Also
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JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
important to the success of this project was the availability of log and core
data, on CD-ROM and at the Web sites of ODP Logging Services and
JANUS. We sincerely thank the Co-Chief Scientists and Shipboard Scientific Party of Leg 185 for making Hole 801C a solidly determined endmember for evolution of upper oceanic crust. We appreciate constructive
reviews by J. C. Alt, C. Broglia, and D. Goldberg. This research used
samples and data provided by ODP. ODP is sponsored by the U.S. National
Science Foundation and participating countries under management of Joint
Oceanographic Institutions, Inc. Funding for this project was provided to
Abrams, Pockalny, and Larson by the U.S. Science Support Program.
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JARRARD ET AL.: PHYSICAL PROPERTIES OF OCEANIC CRUST
Figure 7. Comparison of velocity and potassium logs for
Hole 801C tholeiites. The strong inverse correlation
between these logs indicates that precipitation of potassium-bearing clays and celadonite occurs mainly in the most
porous (lowest-velocity) zones.
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Figure 14. Relationships between porosity and matrix
density based on core index property measurements for the
sites of Table 3. High porosity permits increased fluid flow
and resulting increased alteration, but this effect is only
evident at older sites.
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