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IN PRESS: JOURNAL OF GEOPHYSICAL RESEARCH 1999
COPYRIGHT AMERICAN GEOPHYSICAL UNION
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Faulting and fault scaling on the median valley floor of the transAtlantic geotraverse (TAG) segment, ~26° N on the Mid-Atlantic
Ridge
DelWayne R. Bohnenstiehl1 and Martin C. Kleinrock
Department of Geology, Vanderbilt University, Nashville, Tennessee
Abstract. A quantitative assessment of faulting on the median valley floor of a slow
spreading ridge is accomplished through the analysis of high-resolution DSL-120 sidescan
sonar and coregistered bathymetric data from the TAG segment near 26°N on the MidAtlantic Ridge. At this location, faulting is exposed within a 3-5 km wide ridge-parallel zone
lying asymmetrically on the eastern half of the median valley floor. Mapped faults have a
normal sense of displacement, are < 2 km in length, and accommodate ~1.5% brittle
extension. Evidence of fault linkage within the fault population includes kinked and bent
fault traces in map view, the development of overlapping fault segments or relay ramps, and
the presence of multiple local maxima in the displacement-distance profiles of some faults.
Faults have a slight tendency to dip to the east, or outward relative to the valley axis, and
exhibit little symmetry of fault dip about the axis of the faulted zone or any other ridgeparallel line. Faults exhibit a roughly linear relationship between maximum fault throw and
fault length, with a mean ratio of 0.030 for the population. Regression of length-frequency
data indicates a power-law distribution, with an exponent of 1.64-1.96, demonstrating that
fractal populations can exist in the mid-ocean ridge environment. The fractal nature of this
length-frequency distribution and the ratio of maximum fault throw to fault length differ
significantly from those described previously for populations of larger abyssal hill faults in
the fast spreading environment, where the distribution is exponential and the throw-to-length
ratio is ~5 times lower. These results suggest that the scaling of fault populations in the
mid-ocean ridge setting may vary as a function of spreading rate and/or fault size.
1. Introduction
The role of tectonic processes in modifying the lithosphere at mid-ocean ridges has been greatly
illuminated by first-order morphological and geophysical studies. A more complete characterization of these
processes, however, has been hindered by the present dearth of quantitative information describing the
nature of faulting at various scales and spreading rates. Characterizing the scaling properties of fault
populations, using a combination of sidescan sonar imagery and bathymetric data, affords an opportunity to
better understand mid-ocean ridge faulting and its role in shaping the seafloor. Fault scaling studies within
the terrestrial environment have already contributed significantly to our general understanding of fault
growth and evolution [e.g., Watterson, 1986; Cowie and Scholz, 1992a, b, c; Dawers et al., 1993; Cartwright
et al., 1995, 1996; Dawers and Anders, 1995] and have provided new insights into the factors controlling
regional seismicity [Cowie et al., 1993b; Scholz, 1997]. The integration of fault scaling data into
morphologic studies is also leading to a better appreciation of the importance of faulting in controlling
terrestrial topography [e.g., Anders and Schlische, 1994; Pickering et al., 1999]. Fault scaling studies have
been less extensive in the mid-ocean ridge setting, however, and thus far largely restricted to the abyssal
hills of the fast spreading East Pacific Rise (EPR) [e.g., Cowie et al., 1993a, 1994; Carbotte and Macdonald,
1994].
We present below the first quantitative assessment of the displacement-length and length-frequency
fault scaling relationships in the slow spreading mid-ocean ridge environment. This is accomplished
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through the analysis of high-resolution DSL-120 sidescan sonar imagery and coregistered bathymetric data
from the median valley floor of the trans-Atlantic geotraverse (TAG) spreading segment of the Mid-Atlantic
Ridge (MAR). The TAG segment is well suited for such a study in that it is a relatively representative slow
spreading ridge segment, and an extensive geological and geophysical database already exists [e.g., Rona et
al., 1975; Purdy et al., 1990; Kong et al., 1992; Tivey et al., 1993; Kleinrock and Humphris, 1996a, b]. Our
DSL-120 data provide an opportunity to examine a population of small (length of 0.15-2 km) normal faults
that represent a relatively early stage of faulting (< 0.5 Myr) in the slow spreading environment. In addition
to investigating the scaling properties of this population, we utilize these high-resolution data to
characterize fault-facing direction, fault interaction, and brittle strain at scales previously not possible in this
tectonic setting.
Such information has the potential to greatly improve existing models of seafloor evolution [e.g., Goff and
Jordan, 1988; Malinverno and Gilbert, 1989; Goff et al., 1995] and mid-ocean ridge tectonism [e.g.,
Tapponnier and Francheteau, 1978; Mutter and Karson, 1992; Shaw and Lin, 1993; Tucholke and Lin,
1994; Tucholke et al., 1998] and allows us to address the following issues regarding faulting at mid-ocean
ridges: (1) Does the preference for inward-facing faults observed on the flanks of a slow spreading ridge
[e.g., Macdonald and Luyendyk , 1977; Carbotte and Macdonald, 1990; McAllister and Cann, 1996] initiate
on the median valley floor or within the median valley walls? (2) Under what conditions can mid-ocean ridge
fault populations develop fractal length-frequency distributions? Although most terrestrial fault
populations exhibit a fractal length-frequency distribution [Villemin et al., 1995], previous studies on the
fast spreading EPR [e.g., Cowie et al., 1993a, 1994; Carbotte and Macdonald, 1994] have reported only
exponential distributions. What factors might affect the development of fractal populations in the midocean ridge setting? (3) What is the ratio of maximum fault displacement to fault length (d/L) in the slow
spreading mid-ocean ridge environment? Several studies on the EPR [e.g., Cowie et al., 1993a, 1994;
Carbotte and Macdonald, 1994] have reported d/L ratios at the extreme low end of those reported in the
terrestrial environment [e.g., Schlische et al., 1996]; are these results representative of all faulting on the
ocean floor? (4) How much brittle strain is accommodated by small-scale faulting on the median valley floor?
Constraining the amount of extension accommodated by faults at different stages of development may
prove fundamental in improving our understanding of fault growth and evolution. Previous MAR strain
analyses, however, have focused primarily on the larger more developed faults found farther off axis, where
strains of ~10% are typically reported [e.g., Macdonald and Luyendyk , 1977; Escartin et al., 1999].
1.1 Study Area
The TAG segment of the MAR is the site of one of Earth’s largest mid-ocean ridge hydrothermal vent
systems and, consequently, has been the subject of numerous geological and geophysical investigations
[e.g., Lattimore et al., 1974; Rona et al., 1975; McGregor and Rona, 1975; McGregor et al., 1977; Eberhart
et al., 1988; Zonenshain et al., 1989; Karson and Rona, 1990; Kong et al., 1992; Kleinrock and Humphris,
1996a, b; Rashid and Kleinrock , 1996]. Seafloor spreading rates have been asymmetric during the last 10
million years, with half-spreading rates of 13 mm/yr to the east and 11 mm/yr to the west [Lattimore et al.,
1974; McGregor and Rona, 1975; McGregor et al., 1977]. The TAG segment is ~40 km long, trends northnortheast, and is bounded by nontransform discontinuities to the south and north at ~25°50' N and ~26°15'
N, respectively (Figure 1) [e.g., Sempéré et al., 1990; Purdy et al., 1990]. In map view the median valley floor
has an hourglass shape, narrowing and shoaling to < 3550 m in depth toward the magmatic center of the
segment (at ~26°N), and widening and deepening to > 4500 m toward the segment ends (Figure 1).
In the vicinity of our study (26°06-12' N), a low-relief median valley floor is bounded abruptly on both
sides by major bounding-wall faults (throws > 100 m) that create a steeply dipping slope to the east and a
more terraced slope to the west (Figure 2). The western half of the median valley floor is dominated by an
unfaulted volcanic morphology constructed from linear volcanic ridges and small hummocks (Figure 2)
[Kleinrock and Humphris, 1996b; Rashid and Kleinrock , 1996]. The eastern half of the valley floor, which
contains the TAG active hydrothermal mound, is dissected by a number of small faults and fissures (Figures
2-4). Faulting within this zone, which we will refer to as the fault and fissure zone or FFZ [after Kleinrock
and Humphris, 1996b], is the main focus of our study.
1.2 Length-Frequency Distributions
Terrestrial fault populations typically maintain fractal, or power-law, length-frequency distributions [e.g.,
Okubo and Aki, 1987; Scholz et al., 1993; Villemin et al., 1995; Watterson et al., 1996; Cello, 1997].
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Numerical and scaled-physical models suggest that such distributions are a consequence of the elastic
interactions between simultaneously growing faults [e.g., Sornette et al., 1990, 1993; Cowie et al., 1993b;
Cladouhos and Marrett, 1996; Ackermann and Schlische, 1997; Spyropoulos et al., 1998]. A power-law
length-frequency distribution is of the form N=aL-D ,where N is the number of faults having length ≥ L, a is
a constant reflecting the size of the population, and D is the power-law exponent. Values of D in fault trace
length-frequency studies typically range between 1.1 and 2.0, with the majority of data sets having D values
nearer to the center of this range (1.3-1.8) [cf. Yielding et al., 1996; Cladouhos and Marrett, 1996].
In contrast to most terrestrial fault populations, the length-frequency distribution of abyssal hill faults
along the EPR follows a negative exponential distribution [Cowie et al., 1993a, 1994; Carbotte and
Macdonald, 1994]. Because not all EPR abyssal hill faults grow simultaneously, Cowie et al. [1994] have
suggested that as a population these faults do not interact sufficiently to develop a power-law distribution.
Fractal populations in the mid-ocean ridge environment might therefore be restricted to within the zone of
active fault formation and growth [Edwards et al., 1991; Cowie et al., 1994]. The width of this zone is poorly
constrained; however, it has been suggested on the basis of morphological evidence that fault growth is
restricted to within 5-30 km of the ridge axis on a fast spreading ridge [e.g., Macdonald, 1982; Edwards et
al., 1991; Alexander and Macdonald, 1996] and to within 15-35 km of the ridge axis on a slow spreading
ridge [e.g., Macdonald and Luyendyk , 1977; Macdonald, 1982; Jaroslow, 1996; Searle et al., 1998].
Faulting within the FFZ of the TAG segment lies within ~5 km of the ridge axis (in lithosphere < ~0.5 Ma)
and exhibits a fractal length-frequency distribution, demonstrating that fractal populations can exist on the
median valley floor of a slow spreading ridge.
1.3 Displacement-Length Scaling
A linear relationship between fault length L and maximum displacement d has been reported in numerous
terrestrial field studies [e.g., Elliot, 1976; Opheim and Gudmundsson, 1989; Dawers et al., 1993; Schlische et
al., 1996]. Cowie and Scholz [1992b, c] have proposed a post-yield fracture mechanics model as a theoretical
basis for this relationship. According to this model, the stress across a fault is supported by a length s of
inelastically deforming material within the fault tip. The length of s increases linearly with L, such that the
stress acting on the fault tip is independent of fault length [Cowie and Scholz, 1992b]. This model predicts
a relationship of the form d = γL, where d for normal faults is typically reported as maximum fault throw and γ
is the displacement to length (d/L) ratio of the fault and represents a critical shear strain necessary for fault
propagation [Cowie and Scholz, 1992b]. The value of γ is suggested to be dependent on both the shear
strength of the rock and the remote stress acting on the fault, and should therefore vary as a function of
tectonic environment and rock type [Cowie and Scholz, 1992b, c].
The Cowie and Scholz [1992b] model assumes, however, that faults grow as isolated structures that
nucleate within brittle lithosphere and propagate laterally as they accumulate displacement. In reality, fault
growth is a more complex process in which fault interaction and fault linkage are likely to induce scatter
within all displacement-length data sets [cf. Peacock and Sanderson, 1991; Cartwright et al., 1995, 1996;
Nicol et al., 1996]. For example, following the linkage of two or more faults, the resulting fault may be
initially deficient in displacement relative to length (i.e., it will exhibit a lower d/L ratio) [Cartwright et al.,
1995, 1996]. Conversely, interactions between overlapping, but nonlinking, faults will tend to impede the
propagation of both faults, leading to elevated d/L ratios [cf. Peacock and Sanderson, 1991; Nicol et al.,
1996; Willemse et al., 1996; Cartwright and Mansfield, 1998]. Despite these interactions and the predicted
dependence of the d/L ratio on tectonic environment and rock type a compilation of terrestrial data for both
normal and thrust faults reveals that a roughly linear d/L relationship is preserved across nearly eight orders
of magnitude, with a mean d/L ratio of ~0.03 [Schlische et al., 1996]. Analysis of the d/L scaling relationship
for EPR abyssal hill faults (L = 5-50 km) suggests a linear relationship with a d/L ratio of 0.007-0.004,
systematically less than the terrestrial mean [Cowie et al., 1993a, 1994; Carbotte and Macdonald, 1994].
We show below that the d/L ratio of relatively small normal faults (L= 0.15-2 km) on the median valley floor
of the TAG segment is consistent with the d/L scaling of terrestrial faults and several times larger than that
reported in the fast spreading EPR abyssal hill environment.
2. Methods
2.1 Data Collection and Processing
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The “Deep-TAG” survey aboard the R/V Knorr (leg 142-V) encompassed an area ~10 km x 20 km, covering
the TAG median valley floor and portions of the rift valley walls using the DSL-120 system [Kleinrock and
Humphris, 1996a, b]. The DSL-120 is a 120 kHz phase-difference, split-beam sidescan sonar system,
generating acoustic imagery and coregistered bathymetry over typical swath widths of ~1000 m and ~800 m,
respectively. The system was towed at ~100 m above the seafloor throughout the survey at an azimuth of
023°, parallel to the ridge axis, with line spacings of 300-700 m. Navigation was accomplished through a
system of long-baseline acoustic transponders. This survey strategy provided nearly 100% bathymetric
coverage and >100% sidescan coverage over the entire survey area, with nearly 200% sidescan coverage
(i.e., 100% in each look direction) within the FFZ (Figure 3).
The DSL-120 system incorporates an area of seafloor of ~2 m2 into each individual measurement, with an
across-track sampling dimension of only 0.15-0.33 m [Kleinrock, 1992; Blondel and Murton, 1997]. Given
the small across-track sampling size and structure-parallel survey geometry implemented, the raw
backscatter amplitude of the system is sensitive to scarps having widths of ~0.25 m or less [Kleinrock ,
1992]. Mosaicked sidescan imagery was gridded at 2 m, primarily for computational manageability, but
unmosaicked sonar imagery having a resolution of 0.5 m remained available and was utilized in interpreting
structures (particularly in the tip regions of faults). Comparison of phase information yields bathymetric
data with a vertical resolution of ~1-2 m. The bathymetric data, somewhat noisier than the backscatter data,
were gridded at 5 m, and measurements of maximum throw are accurate to within ~2-5 m.
2.2 Fault Identification and Analysis
Fault traces within the survey area were mapped by combined analysis of both sidescan sonar imagery
and bathymetric data at scales of 1:10,000 and greater. Fault scarps can be distinguished from volcanic and
mass-wasting features by the steepness of their slopes in this high-resolution bathymetric data set and by
the crisp, linear, high-amplitude nature of their acoustic backscatter in the sidescan imagery, as illustrated in
Figure 5 [Kleinrock, 1992; Blondel and Murton, 1997]. Fault length was measured as the straight line
distance between fault tips in the sidescan imagery (i.e., A-A’ in Figure 5). Two faults were considered
linked if their respective fault traces were parallel to each other and separated by < 15 m in the sidescan
imagery. Although fault traces separated by as little as 2 pixels (4 m) can be distinguished in the mosaicked
sidescan imagery, we established the 15 m criterion to allow for a more direct comparison between the data
collected with the DSL-120 system used at TAG and the data collected with the coarser resolution GLORIA
and SeaMARC II systems used on the EPR flanks (see Sections 5.2 and 5.3). The latter systems,
unfortunately, cannot distinguish fault traces separated by less than ~150 m [Kleinrock , 1992; Cowie et al.,
1994; Blondel and Murton, 1997]. Since the scale of fault segmentation within the FFZ (segment lengths <1
km) is an order of magnitude less than that observed on the EPR (segment lengths of 1-10 km and greater),
this 15 m criterion minimizes sampling bias between data sets by proportionally scaling the offset criterion
used to establish linkage with the scale of fault segmentation. A linear scaling is justified because the level
of elastic interaction between two faults scales linearly with the length of both faults [Pollard and Segall,
1987; Cladouhos and Marrett, 1996; Cowie and Shipton, 1998]. Two 5 km long fault segments offset by 150
m in the data set of Cowie et al. [1993a] are therefore roughly equivalent to two 0.5 km long fault segments
separated by 15 m in our data set.
Errors associated with measuring fault length arise primarily from the finite resolution of the data. As the
tip of a fault is approached, the offset on the scarp will decrease below the resolution of the bathymetric
data first and then the sidescan imagery. To minimize the under-sampling of fault lengths, all fault length
measurements were obtained from the sidescan imagery. The ability of a sonar system to image a fault scarp
begins to fall off when the offset on the scarp falls below some fraction of the across-track footprint size or,
in the case of gridded data, the pixel size [Blondel and Murton, 1997]. For the DSL-120 data used in our
study, scarps with offset < 0.5 m and < 2 m may fail to be fully represented in the unmosaicked and
mosaicked data, respectively. For the purpose of later error analysis, we will assume that portions of a fault
tip with throw < 2 m are not properly imaged. We emphasize, however, that this represents a maximum
unresolved throw at the mapped fault tip, since only some portion of the sonar footprint need insonify the
fault scarp to generate significantly higher backscatter, relative to the unfaulted terrain. The GLORIA and
SeaMARC II systems used on the EPR may fail to image portions of a fault scarp having < 30-45 m of throw
[Cowie et al., 1994], and therefore a significantly greater length of each fault fails to be properly imaged. To
the first order, however, the proportionality which exists between survey resolution and the scale of faulting
investigated in the TAG and EPR studies serves to equilibrate any potential sampling bias associated with
instrument resolution (i.e., the percentage of total fault length not imaged remains largely unchanged). We
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investigate how data resolution influences the length-frequency distribution of both data sets below (see
Sections 3.1 and 5.2).
Maximum throw was not determined from the scarp height at the center of the fault but rather was
selected from a series of bathymetric profiles (extracted from the gridded data) crossing perpendicular to the
fault (Figure 5). Displacement data generated in this manner assume the effects of sedimentation and
volcanic infilling are minimal. We avoided measuring fault throws in areas where these assumptions are not
valid (e.g., along faults at the edge of the neovolcanic zone and in areas obscured by submarine landslides).
Since bathymetric coverage is < 200% (i.e., < 100% in both look directions), we have restricted our throw
measurements to fault scarps that were insonified, avoiding areas where it would be necessary to estimate
throw in acoustic shadows.
3. Faulting on the Median Valley Floor
3.1 Fault Distribution and Scaling
We mapped 200 normal fault traces, ranging in length from 150 m to 2050 m, within the FFZ of the TAG
median valley floor (Figure 4). These faults have a slight (65%) tendency to dip to the east, or outward
relative to the ridge axis, and maintain a mean strike of 027°, slightly clockwise of the 023° trend of the
spreading segment. We find no positive correlation between dip direction and orientation, and dip direction
shows little symmetry about the axis of the FFZ or any other ridge-parallel line (Figures 4 and 6). Fault
spacings range from tens to hundreds of meters, in contrast with much larger spacings (4-15 km) typical in
off-axis MAR abyssal hill fault populations [e.g., Shaw and Lin, 1993; Jaroslow, 1996].
Figure 7a illustrates the roughly linear relationship between maximum fault displacement (reported as
maximum throw d) and fault length L within the FFZ population. The d/L ratio (γ) of individual faults ranges
from 0.017 to 0.085, with a mean value of 0.030 for the population (consistent with the terrestrial mean [e.g.,
Schlische et al., 1996]). Similar degrees of scatter are common in most d/L data sets and may be attributed to
fault linkage and fault interaction [cf. Cartwright et al., 1995,1996].
The linear trend in the log-log length versus cumulative frequency plot (Figure 7b) reveals FFZ faults
follow a fractal distribution over nearly an order of magnitude, with a power-law exponent D of 1.64.
Truncation at the longest lengths may result from the longest faults extending out of the survey area [e.g.,
Gillespie et al., 1993; Yielding et al., 1996] or potentially from portions of the longest faults being obscured
by the large landslide that dominates the morphology of the northern FFZ and effectively limits the survey
area (Figure 4). Truncation at shorter lengths in Figure 7b results from failure of the DSL-120 to resolve
some portion of the fault tip [e.g., Gillespie et al., 1993; Yielding et al., 1996]. The percentage of the total
fault length not imaged is greater for a small fault relative to a larger fault; therefore, this bias in sampling
preferentially affects the shortest faults in the population.
The undersampling of the true fault length due to the limited resolution of the survey instrumentation can
significantly influence the estimated value of D [Yielding et al., 1996; Pickering et al., 1997]. This may be
corrected for by estimating the throw at the imaged or mapped fault tip and then extending the fault by
assuming some d/L ratio [e.g., Pickering et al., 1997]. We will adopt a value of 2 m, which should be
considered a maximum, for the throw at the mapped tip of each fault (see Section 2.2). Assuming our
observed d/L ratio of 0.030, as much as 68 m should be added to the raw length measurements (i.e., 34 m to
each tip) to account for the failure of the DSL-120 to fully image the fault tips. Figure 7c shows the corrected
length-frequency plots for the “L+ 68” data. These data are also consistent with a fractal distribution (i.e.,
the fundamental nature of the distribution has not changed); however, the estimated value of D has
increased to 1.92. Recalculating the d/L ratio of the population using these adjusted L+68 data reduces its
value only slightly to 0.029.
3.2 Estimating Strain
Our data provide an opportunity to estimate the amount of extensional brittle strain accommodated by a
population of small faults on the median valley floor of a slow spreading ridge. Placing constraints on the
amount of extension accommodated by populations of mid-ocean ridge faults in different stages of
development is fundamental to improving our understanding of fault growth in this tectonic setting.
Previous strain analyses along the MAR indicate strains of ~10% are reached within the crests of the
median valley [e.g., Macdonald and Luyendyk , 1977; Jaroslow, 1996; Escartin et al., 1999]. Our high-
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resolution DSL-120 data allow us to examine an earlier stage of fault development on the median valley floor
prior to the onset of bounding-wall faulting.
We sum the extension, or heave H, along all faults within the FFZ and divide by the total faulted area A to
obtain an estimate of the extensional two-dimensional (2-D) brittle strain εx on the upper surface of this
region: εx = A-1∑H. The average heave on a fault is γavg L/tan θ (where θ is fault dip and γavg is mean throw),
or γL/2 tan θ, assuming an inelastic model of fault displacement where the mean throw is equal to half the
maximum throw on a fault [Scholz and Cowie, 1990]. The total extension contributed by a single fault is
therefore εx = γL2/2A tan θ; summing this for all faults provides the 2-D strain. We choose to apply this 2-D
method over a 1-D summation of displacement [e.g., Macdonald and Luyendyk , 1977; Kleinrock and Hey,
1989; Escartin et al., 1999] because it avoids the necessity of estimating throw within acoustic shadows,
allows us to provide additional constraints on the extension accommodated by faults below the resolution
of our data (as described below), and is similar to the methods used to estimate strain in a number of field
and laboratory studies that we refer to in our later discussion [e.g., Cowie et al., 1993a; Spyropoulos et al.,
1999]. Moreover, this 2-D method of strain estimation allows us to easily account for (and exclude) areas
within the FFZ that have been overprinted by volcanism and mass-wasting (Figure 4).
Analysis of DSL-120 bathymetric profiles (e.g., Figure 5) indicates that fault-line scarps within the FFZ
maintain slopes up to ~60°. A 60° fault dip is consistent with that reported for focal mechanism solutions
associated with much larger normal faults in the TAG eastern bounding wall [e.g., Kong et al., 1992], as well
as with Andersonian theory [Anderson, 1951]. However, fault dips measured from scarp slopes are likely to
represent an underestimation of the true fault dip because of the spatial averaging that occurs when the
bathymetry data are gridded and the effects of mass wasting on the fault scarp. Moreover, Quaternary fault
scarps in terrestrial rifts, such as Iceland, typically exhibit near-vertical dips at the surface [e.g.,
Gudmundsson, 1987a,b], and similar size normal faults in the Volcanic Tablelands of eastern California
exhibit dips of ~70° [Dawers et al., 1993]. Thus, a reasonable estimate of strain can be made by assuming
fault dips of 75° ±15°.
Excluding areas obscured by landslides or volcanism, the total faulted area A of the FFZ is estimated to be
~15 km2. Taking γ = 0.030 and θ = 75°, we estimate that a brittle strain of ~1.2% is accommodated by faulting
within the FFZ (a range of 0-2% strain is possible if we consider θ from 90° to 60°). Analyzing the “L+68”
corrected data (see Section 3.1) increases this by ~17%. However, these estimates ignore the strain
potentially contributed by faulting at scales less than the resolution of our sidescan data (i.e., fault lengths <
150 m). We may estimate the amount of strain not resolved by the DSL-120 data by extrapolating the
observed power-law length-frequency distribution [cf. Scholz and Cowie, 1990; Villemin et al., 1995].
Integrating the product of the probability density function for a power-law distribution (-dN /dL) and our
estimate of εx for a single fault, the total strain contributed by faults below the resolution of our data can be
expressed as εx = γaDLmin D-2/2(2-D)A tan θ, where Lmin is the length of the smallest mapped fault (i.e., 150 m in
the raw length data). This analysis suggests that an additional strain of ~0.6% may be accommodated by
faults < 150 m in length. Synthesizing these results, we estimate the strain associated with FFZ faults on the
median valley floor to be ~1.5 ± 1%, nearly an order of magnitude less than the strain reported throughout
the broad rift valley system [e.g., Macdonald and Luyendyk , 1977; Escartin et al., 1999].
3.3 Fault Linkage and Fault Interaction
Fault linkage and interaction have been documented extensively in many terrestrial fault population
studies [e.g., Peacock and Sanderson, 1991, 1994; Cartwright et al., 1995; Dawers and Anders, 1995;
Cartwright and Mansfield, 1998]. Sonar surveys suggest that fault linkage is also an important mechanism
in the growth of abyssal hill faults at all spreading rates [e.g., Cowie et al., 1994; McAllister and Cann,
1996; Searle et al., 1998; Shaw et al., 1998]. Although fault linkage occurs at all scales, it has been
suggested that fault linkage becomes increasingly important, and fault nucleation becomes less important,
as a fault population evolves [e.g., Cowie et al., 1995; Wojtal, 1996; Spyropoulos et al., 1998, 1999].
We observe some evidence of fault linkage and interaction at the scale of median valley floor faulting,
within a population of faults that accommodates significantly less strain and is presumably less evolved
than those studied farther off axis [e.g., McAllister and Cann, 1996; Searle et al., 1998]. We summarize this
evidence in Figure 8 by examining the region directly to the east of the TAG hydrothermal mound. Fault
linkage and interaction are particularly well developed and exposed in this region, as is evident by (1) kinked
and bent fault traces in map view, which may indicate the physical linkage of faults that did not propagate
directly into each other along strike (Figures 8b and 8c) [e.g., Peacock and Sanderson, 1994]; (2) regions of
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fault overlap or soft linkage, often referred to as relay ramps, in which more obliquely oriented structures
may develop to accommodate linkage between fault segments (Figures 8b and 8c) [e.g., Larsen, 1988;
Peacock and Sanderson, 1991, 1994]. McAllister and Cann [1996] have described obliquely oriented
structures, similar to the ones shown in Figure 8, connecting overlapping fault segments separated by as
much as 1.5 km in the bounding walls of the MAR; and (3) multiple local maxima in the displacement-length
profiles of some faults (Figure 8a, segment D), with the intervening point(s) of minimum displacement
possibly indicating a point of fault linkage [e.g., Peacock and Sanderson, 1991].
Because faults cannot propagate into the stress shadow of a neighboring fault, the overlap or soft linkage
of faults tends to impede their propagation. Fault interaction should therefore lead to elevated d/L ratios
and steeper displacement gradients within the tip regions of faults. These effects have been documented
extensively in the terrestrial environment [e.g., Peacock and Sanderson, 1991; Nicol et al., 1996; Willemse et
al., 1996] and potentially observed within the abyssal hill faults of the EPR, where deep-towed sonar
observations indicate that fault segments may maintain d/L ratios as much as ~2-3 times larger than that
associated with the entire fault [Cowie et al., 1994]. Such effects also appear evident in segment A, which is
overlapped at both its tips. Segment A maintains a d/L ratio of 0.043 (above the 0.030 mean) and has a steep
displacement gradient at both tips relative to the southern tip of fault 2, which is not overlapped by any fault
of significant size. The linkage of fault segments, on the other hand, will result in initially lower d/L ratios
[e.g., Cartwright et al., 1995, 1996]. This is evident in fault 1, which consists of four segments and
maintains a d/L ratio of 0.022 (below the 0.030 mean).
It has been suggested that the cumulative displacement profile of faults or fault segments that are
strongly interacting (but not physically linked) will mimic that of an isolated fault [e.g., Dawers and Anders,
1995; Cowie, 1998]. The dashed line in Figure 8a represents the summation of throw along faults 1 and 2,
which are offset by ~ 45 m at their closest point. The cumulative profile for both faults does not represent
that expected for an isolated fault (i.e., a roughly elliptical to bell-shaped profile [e.g., Barnett et al., 1987;
Dawers et al., 1993; Schlische et al., 1996]), suggesting that the entire structure lacks sufficient interaction
to be considered a single fault. Interpreting faults 1 and 2 as a single fault, however, would yield a d/L ratio
of 0.015, consistent with the d/L scaling observed for fault segments comprising the abyssal hill faults
described by Cowie et al. [1994] on the EPR. The fault segments described by Cowie et al. [1994] do not,
however, appear themselves to be noticeably segmented within the resolution of the SeaMARC I sonar data
used in that study (~2.5 m pixels). The displacement profile of fault 1 also departs somewhat from that
expected for an isolated fault, largely because of a minimum in displacement associated with what we
interpret to be a point of linkage within segment D.
4. Median Valley Floor/Bounding Wall Transition
The transition from the median valley floor faults described above to the much larger bounding
wall/abyssal hill faults appears to be abrupt, with many of the details of this transition unfortunately
obscured by mass-wasted debris from the valley walls. Although we imaged the first several bounding wall
faults, the limits of the survey preclude any statistically significant determination of the scaling properties of
this population. Nonetheless, there exist three first-order differences between the neighboring valley floor
and bounding wall fault populations: (1) bounding wall faults increase in length (up to ~8 km, limited by the
survey dimensions), throw (~100-300 m), and spacing (~1-3 km); (2) the percent of faults dipping away from
the rift axis decreases to < 10% within the bounding walls; and (3) while we would expect small faults similar
to those imaged on the valley floor to be rafted off between larger bounding wall faults, these structures are
extremely scarce even in our high-resolution DSL-120 data. This is likely a consequence of pronounced
mass wasting on the valley walls, which manifests itself both in localized landslides, resembling terrestrial
slump deposits, and more diffuse down-slope transport [e.g., Getsiv et al., 1998].
5. Discussion
5.1 Inward Versus Outward Dipping Faults
The percentage of inward dipping abyssal hill faults has been shown [e.g., Carbotte and Macdonald,
1990] to vary as a function of spreading rate, with faults on the flanks of slow spreading ridges being
typically > 80% inward dipping and faults on the flanks of fast spreading ridges being ~50% inward
dipping. Searle [1984] and Carbotte and Macdonald [1990, 1994] have suggested that the dominance of
8
inward dipping faults in the slow spreading environment arises because inward dipping abyssal hill faults
will extend to much shallower depths than their outward dipping counterparts before becoming constrained
at depth by the brittle-plastic transition (a consequence of the rapid thickening of the brittle layer off axis).
This decreases the mean normal stress acting on inward dipping faults, relative to outward dipping faults,
and implies that inward dipping faults will grow preferentially in the slow spreading environment. Along
fast spreading ridges the thickness of the brittle layer increases more gradually off axis. This effect is
therefore minimized, giving rise to a more equal abundance of inward and outward dipping faults [Carbotte
and Macdonald, 1990, 1994].
In the above model, fault growth is only affected when inward dipping faults reach a sufficient size to
intersect the brittle-plastic transition. Seismic data [Kong et al., 1992] and thermal models [Shaw and Lin,
1996] suggest the brittle-plastic transition occurs at a subseafloor depth of 3-5 km beneath the TAG FFZ.
Given that faults are unlikely to extend to depths greater than their length [Nur, 1982; Gudmundsson, 1987a,
b], the length distribution shown in Figure 7b implies that TAG median valley floor faults do not intersect
the brittle-plastic transition. The more equal distribution of inward and outward dipping faults observed in
the TAG FFZ population may therefore be attributed to the Carbotte and Macdonald [1990, 1994] model
because neither dip direction preferentially reaches the brittle-plastic transition, and hence there should exist
no systematic difference in mean normal stress acting on these faults.
Within the bounding walls, however, both the greater size of faults [Jaroslow, 1996; McAllister and
Cann, 1996; Searle et al., 1998], as well as observed hypocenter depths and focal mechanisms [Kong et al.,
1992; Wolfe et al., 1995], implies that inward dipping bounding wall faults may be constrained at the depth of
the brittle-plastic transition [cf. Harper, 1985]. Under these conditions the mechanism described by
Carbotte and Macdonald [1990, 1994] can work to limit the further growth of outward dipping faults,
perhaps explaining the low-abundance of these structures observed within the bounding walls of the TAG
and other segments of the MAR [e.g., Macdonald and Luyendyk , 1977; McAllister and Cann, 1996; Searle
et al., 1998; Escartin et al., 1999].
5.2 Length-Frequency Distributions
Theoretical and experimental evidence suggests that two conditions must be met to facilitate the
formation of a fractal distribution in a population of faults: (1) both long-and short-range elastic interactions
must occur between faults in the population [Sornette et al., 1990, 1993], and (2) fault growth must remain
nucleation-dominated rather than linkage-dominated [e.g., Spyropoulos et al., 1998; 1999]. Violating either
of these conditions has been shown to give rise to exponential distributions [e.g., Malinverno and Cowie,
1993; Ackermann and Schlische, 1997; Spyropoulos et al., 1998; 1999]. Condition 1 should be met provided
faults within the population grow simultaneously with one another [e.g., Cowie et al., 1994] and they do not
extend completely through the brittle layer [Ackermann and Schlische, 1997]. Condition 2 should be met
within populations of faults accommodating low (< ~2.5%) strains [Spyropoulos et al., 1998, 1999].
The fractal nature of the TAG FFZ population (Figure 7) shows that in accordance with these conditions
a population of simultaneously active mid-ocean ridge faults (within < 5 km of the ridge axis in lithosphere <
~0.5 Ma) accommodating ~ 1.5% strain and having lengths less than the thickness of the brittle can exhibit a
fractal length-frequency distribution. This power-law length-frequency scaling is consistent with most
terrestrial fault populations [e.g., Villemin et al., 1995], but contrasts with the exponential distributions
previously reported on the EPR [e.g., Cowie et al., 1993a; 1994; Carbotte and Macdonald, 1994]. To verify
that the exponential nature of faulting on the EPR is real and cannot be attributed to survey resolution, we
applied the method outlined by Pickering et al. [1997] (as described above) to the raw EPR lengthfrequency data of Cowie et al. [1994] (one GLORIA and one SeaMARC II data set). These analyses show
that even after accounting for limited sonar system resolution, the data cannot be satisfactorily fit with a
power-law model and are fit substantially better by an exponential model. Hence the exponential nature of
faulting on the EPR appears genuine.
Several conditions consistent with the development of an exponential, rather than fractal, distribution
exist within the previously studied EPR fault population. (1) The EPR studies covered a region extending
well outside the active zone of fault growth, limiting the degree of interaction within the population [Cowie
et al., 1994]. (2) The lengths of the faults studied are of sufficient size (mean length 6-10 km [Cowie et al.,
1994]) to penetrate the thin (~1-3 km thick) brittle lithosphere that exists in the fast spreading environment.
Scaled-physical models of extensional environments have shown that the penetration of the largest faults in
a population through a mechanical layer leads to the suppression of long-range elastic interactions and the
9
development of exponential distributions [Ackermann and Schlische, 1997]. Since faults in the fast and
superfast spreading environments may penetrate the thin brittle layer early in their development, this
mechanism suggests that fractal distributions may be extremely rare in these settings. (3) Abyssal hill faults
accommodate strains on the order of 5-15% [e.g., Cowie et al., 1993; Carbotte and Macdonald, 1994].
Scaled-physical and numerical models suggest that fractal length-frequency distributions will begin to break
down in populations of faults accommodating > 2.5%, because of the increasingly important role of fault
linkage [Spyropoulos et al., 1998, 1999]. These models suggest a transition to fully exponential distributions
at strains of 10-15%.
We have shown that fractal populations can exist at the scale of median valley floor faulting in the slow
spreading environment; however, more work is needed to determine across what range of scales and
spreading rates fractal populations might exist in the mid-ocean ridge setting. It seems unlikely that the
entire zone of active fault growth will maintain a fractal distribution, because faults on the outer edge of this
zone will have undergone considerable growth before the nucleation of the youngest faults within the zone.
Rather, fractal populations may be limited to smaller fault packets [Cowie et al., 1994] that nucleate
penecontemporaneously and continue to grow simultaneously. Such packets might be initiated during a
sequence of diking events, similar to those observed in Iceland, where fractal distributions have been
reported over a scale range comparable to that observed within the FFZ [e.g., Gudmundsson, 1987a,b].
5.3 Displacement-Length Ratios
We have shown that small (L < 2 km) faults on the median valley floor of the slow spreading TAG ridge
segment maintain d/L ratios of ~0.030, consistent with the scaling of terrestrial fault populations. Previous
work, with coarser resolution sonar systems on the EPR, suggests that larger (L ~5-50 km) faults in the fast
spreading environment maintain d/L ratios several times lower (0.007-0.004). Clearly, care must be taken in
interpreting displacement-length data collected by instruments of differing resolution, because lowerresolution data sets will be biased toward imaging larger structures. This could potentially lead to the
misinterpretation of several mechanically separate faults as a single fault, resulting in a lower estimated d/L
ratio [cf. Cowie et al., 1994].
To minimize the effects of possible fault misinterpretation in the SeaMARC II and GLORIA data sets, we
have adjusted the offset criterion used to establish linkage in the DSL-120 data set so that it is scaled
proportionally to that criterion used on the EPR (see Section 2.2). Despite doing so, TAG valley floor faults
do not exhibit the extremely low values of γ reported on the EPR. This suggests that the d/L scaling of EPR
abyssal hill faults is the result of some physical process rather than sampling bias and hints at the existence
of a spreading-rate and/or scale dependence of the d/L ratio in the mid-ocean ridge environment. Such a
dependence might potentially be linked to a number of factors, such as variations in the shear strength of
the lithosphere [e.g., Cowie and Scholz, 1992b, c], the relative role of fault linkage at different scales and
spreading rates [e.g., Spyropoulos et al., 1999], the limited throw that each fault can obtain before locking
[e.g., Shaw and Lin, 1993], the breaking of faults through a brittle layer of variable thickness [e.g., Dawers et
al., 1993; Shaw and Lin, 1996], and the initiation of faults as tensile fractures [Gudmundsson, 1992]. Future
displacement-length analyses off axis on the TAG segment, using data of equivalent resolution to that used
previously on the EPR, will help to clarify the effects of survey resolution and elucidate what factors control
the observed variation in the d/L ratio.
6. Summary and Conclusions
Analysis of a high-resolution DSL-120 sidescan sonar and coregistered bathymetry data set, collected
near ~26°N on the MAR, has allowed us to investigate the nature of faulting within the median valley of a
slow spreading ridge at scales previously unavailable. This paper represents the first quantification of the
length-frequency distribution and displacement-length ratio of a fault population in the slow spreading midocean ridge environment. We summarize below our observations regarding the distribution and scaling of
the relatively small (0.15-2.0 km) faults in the fault and fissure zone (FFZ) of the TAG median valley floor.
Many of these observations may be applicable to some other slow spreading median valleys as well.
1. Faulting on the median valley floor of the TAG segment is exposed within a 3-5 km wide ridge-parallel
zone that lies asymmetrically on the eastern half of the valley floor. Mapped faults appear to have a normal
sense of displacement, maintain predominantly ridge-parallel orientations, and are < 2 km in length.
10
2. Estimates of brittle extension suggest that these faults accommodate ~ 1.5 ± 1% brittle extension
(assuming fault dips of 60°-90°), nearly an order of magnitude less than that accommodated beyond the
valley floor on the TAG [e.g., Jaroslow, 1996] and other MAR segments [e.g., Macdonald and Luyendyk ,
1977; Escartin et al., 1999].
3. Evidence of fault linkage on the valley floor includes kinked and bent fault traces in map view, the
development of overlapping fault segments or relay zones, and the presence of multiple local maxima in the
displacement-distance profiles of some faults. These observations indicate that even at the scale of median
valley floor faulting, fault growth is to some degree accomplished by the linkage of fault segments. The
amount of strain accommodated by the FFZ population (< 2.5%), however, suggests that the primary
mechanism of fault growth within the population has been nucleation rather than linkage [cf. Spyropoulos
et al., 1999]. In contrast, bounding wall fault growth is probably more linkage-dominated, as suggested by
extensive morphological evidence [e.g., McAllister and Cann, 1996; Searle et al., 1998] and the greater
amount of strain (>10% [e.g., Jaroslow, 1996; Escartin et al., 1999]) accommodated within the bounding
walls.
4. Inward and outward dipping faults are present in near equal abundance at the scale of median valley floor
faulting, in contrast to the dominance (> 90 %) of inward dipping faults mapped within the bounding walls
and off axis on the MAR. We suggest that TAG median valley floor faults are too shallow to intersect the
brittle-plastic transition, and therefore the mean normal stress mechanism proposed to explain the
dominance of inward dipping faults off axis [e.g., Carbotte and Macdonald, 1990] does not influence the dip
direction of slow spreading faults at this scale.
5. Faults maintain a fractal length-frequency distribution with a power-law exponent of ~ 1.64-1.92. These
data show that a fractal population of faults can exist within the zone of active fault growth along a midocean ridge (within ~5 km of a slow spreading ridge axis and in lithosphere < 0.5 Ma), provided faults do not
penetrate the brittle layer and accommodate strains consistent with a nucleation-dominant pattern of
growth.
6. The maximum throw versus length (d/L) ratio of faults on the median valley floor of the TAG segment
(0.030) is consistent with that reported for faulting across a range of scales in the terrestrial environment but
is a factor of 4-7 times greater than that reported for larger abyssal hill faults on the flanks of the fast
spreading EPR. These data suggest that the d/L ratio of faults in the mid-ocean ridge environment may vary
as a function of spreading rate and/or scale; however, further work across a broader range of scales and
spreading rates is necessary to clarify what factors control the observed variation in the d/L ratio.
Acknowledgments. We thank the captain, officers, and crew of the R/V Knorr, as well as members of the
science party on Leg 142-V, for their excellent work. Particular gratitude is extended to the Woods Hole
Oceanographic Institution's Deep Submergence Operations Group for their highly successful operation of
the DSL-120 system. We especially thank co-chief scientist Susan Humphris for her hard work, dedication,
wisdom, and fine attitude that were so crucial to the overall success of this "Deep-TAG" cruise. We also
thank J. Lin, B. Tucholke, A. Gupta, and P. Shaw for interesting discussions about faulting processes; C.
Spyropoulos, N. Dawers, J. Contreras, and C. Scholz for making various manuscripts available to us, R.
Buck, S. Carbotte, A. Clifton, A. Hosford, and C. Scholz for further discussion and reviewing various drafts
of this manuscript, P. Cowie for making L-frequency data from the EPR available to us, M. Rashid and T.
Larrieu for technical assistance; J. Nagel for proofing this manuscript; P. Wessel and W. Smith for
development of GMT; and D. Caress and D. Chayes for development of the MB-system. Reviews by D.
Wright and C. Ebinger greatly improved the quality of this manuscript. This work was supported by NSF
grants OCE9314697 and OCE9712230.
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_________
D. Bohnenstiehl, Lamont-Doherty Earth Observatory, Box 1000, Palisades,
NY 10964. ([email protected])
M. Kleinrock, Department of Geology, Vanderbilt University, Box 1805
Station B, Nashville, TN 37235. ([email protected])
(Received July 24, 1998; revised June 10, 1999;
accepted July 14, 1999)
________
1
Now at Lamont-Doherty Earth Observatory, Department of Earth and
Environmental Sciences, Columbia University, Palisades, NY.
Copyright 1999 by the American Geophysical Union.
Paper number 1999JB900256.
0148-0227/99/1999JB900256$09.00
Figure Captions
Figure 1. Combined Sea Beam, Hydrosweep, and DSL-120 bathymetric map of the TAG segment of the
Mid-Atlantic Ridge, from the data of Purdy et al. [1990], Tucholke et al. [1997], and Kleinrock and
Humphris [1996a], respectively. Solid line show the area covered by the 1994 DSL-120 survey. Dashed
lines show the axes of the spreading segments [e.g., Semp éré et al., 1990]. Contour interval is 200 m.
Figure 2. Geologic interpretation and data examples from the TAG segment of the MAR. (a) Crosssection NW-SE crossing near the TAG active mound (star). Vertical exaggeration is 3 times. (b) Data
summary and interpretation map of the ~10 x 20 km region imaged by the DSL-120 system. Kleinrock and
Humphris [1996b] have divided the median valley into four morphologic zones (from west to east: western
bounding wall zone, neovolcanic zone, fault and fissure zone (FFZ), and eastern bounding wall zone). The
V-pattern in the neovolcanic zone represents most recent volcanics. Thin lines represent structures
within the FFZ. Thick solid lines represent bounding wall faults. Ticks are shown on down-thrown
blocks. (c-f) Characteristic sidescan data examples from the four zones defined by Kleinrock and
14
Humphris [1996b]: (c) western bounding wall zone, (d) neovolcanic zone, (e) FFZ, and (f) eastern
bounding wall zone. Strong returns are light. Acoustic shadows are black. FFZ is insonified from the
east, other zones are insonified from nadir (artifacts running vertically up image). See Figure 4 for
distribution of mapped faults within the FFZ. After Kleinrock and Humphris [1996b].
Figure 3. DSL-120 sidescan imagery of area including the FFZ. Light shading represents high amplitude
backscatter. Tracklines trend 023°. (a) “East-looking” mosaic of DSL-120 sidescan imagery. Data
insonified from the west-northwest are preferentially overlain. b) “West-looking” mosaic of DSL-120
sidescan imagery. Data insonified from the east-southeast are preferentially overlain. See Figures 5 and 8
for more detailed images.
Figure 4. Interpretive map of the area shown in Figure 3. Thin shaded lines represent outward or east
dipping fault traces. Thin solid lines represent inward or west dipping faults. The morphology of the
northern portion of the FFZ is dominated by debris from a large landslide that originated in eastern
bounding wall. The star shows the location of the TAG active hydrothermal mound.
Figure 5. Illustration of how fault lengths and throws were determined from DSL-120 data. (a) DSL-120
sidescan imagery. High-amplitude returns are light. Insonification from the east-southeast. Linear region
of high-amplitude backscatter (A-A’) represents the trace of an east-dipping fault. (b) Sidescan intensity
and bathymetry across the profile B-B’. Fault scarps are distinguishable because of their high amplitude
backscatter and steep slopes and their crisp linear character in map view.
Figure 6. Histogram of fault strike. Faults have a mean strike of 027°, slightly east of ridge parallel for the
segment (023°).
Figure 7. a) Maximum fault throw versus fault length for faults on the median valley floor of the TAG
segment, as determined from DSL-120 bathymetric data and sonar imagery, respectively. Vertical error bars
±5 m and horizontal error bars 0 to +68 m to account for system resolution. Maximum throw versus length
ratio of EPR abyssal hill faults is shown for reference. (b) Log-log cumulative frequency plot of mapped
fault trace lengths on the median valley floor of the TAG segment. The vertical axis is the number of
faults having lengths greater than or equal to the corresponding fault length, shown on the horizontal
axis. The data display a fractal length-frequency distribution with a power-law exponent of 1.64. (c) Loglog cumulative frequency plot of corrected fault trace lengths (L + 68 m) on the median valley floor of the
TAG segment. The data display a fractal length-frequency distribution with a power-law exponent of 1.92.
Figure 8. Examples of fault linkage and interaction within the FFZ. (a) Throw-distance profiles for the
faults shown in Figures 8b and 8c. Profiles were generated from multiple bathymetric profiles crossing
perpendicular to the fault array, as shown in Figure 5. Segment A (circles), segment C (diamonds),
segment D (crosses), fault 2 (hexagons), cumulative profile (dashed line). Bathymetry was not obtainable
along portions of segment B, and therefore its profile is not included. See text for discussion. (b) Fault
traces mapped from image shown in Figure 8c and discussed in the text. Ticks are marked on down
thrown block. (c) DSL-120 sidescan imagery from which fault traces were mapped. High-amplitude
backscatter returns are light. Nadir is evident as a horizontal artifact in the image. All faults mapped in
Figure 8b are illuminated from the west (i.e., the bottom of Figure 8c).
Data File sbeam.tagseg.100m.sur.grd
o
26 20 ' N
o
26 10 ' N
o
26 00 ' N
o
25 50 ' N
o
45 10' W
o
45 00' W
o
44 50' W
-5000 -4700 -4400 -4100 -3800 -3500 -3200 -2900 -2600 -2300 -2000
Depth (m)
Figure 1: Bohnenstiehl and Kleinrock "TAG valley floor faulting"
Eastern Bounding
Wall Zone
2500 m
Neovolcanic Fault and
Fissure
Zone
Zone
(FFZ)
Western Bounding
Wall Zone
3500 m
300 o
V VV V V V
*
V
120 o
'
'N
10
26
12
'
'
48
'
50
52
'
V VV V
14
a)
54
'W
FFZ
Youngest
Volcanics
1
8
26
4'N
'N
6'
'N
'N
12
26
8'
N
50
'W
'W
52
b)
26
'N
5
06
km
44
0
'W
'W
48
54
44
44
10
26
e)
c)
d)
500 m
f)
Figure 2: Bohnenstiehl and Kleinrock "TAG valley floor faulting"
0
0
44 48' N
44 48' N
1 km
1 km
km
1
N
NN
0
26 10' N
0
44 50' N
0
26 10' N
0
44 50' N
0
26 10' N
26 10' N
0
0
44 48' N
0
44 48' N
26 08' N
0
26 08' N
0
26 08' N
0
0
26 08' N
b)
a)
0
44 50' N
0
44 50' N
ure 3a: Bohnenstiehl and Kleinrock "TAG valley floor faulting"
Figure 3b: Bohnenstiehl and Kleinrock "TAG valley floor faulting"
Figure 3a,b Bohnensteihl and Kleinrock: "TAG valley floor faulting"
0
44 48' N
West
Dipping Faults
East
Dipping Faults
1 km
N
0
26 10' N
Vo
lca
nic
Te
rra
in
Northern
Boundary of Survey
0
44 50' N
NeoVolcanic
Zone
26 10' N
44 48' N
lca
nic
0
Te
0
rra
in
Landslide
Deposit
M
from ass Was
t
East
ern B ed Debri
s
ound
ing W
all
anic
Volc
hern
Boun
dary
0
0
26 08' N
Sout
26 08' N
Terra
in
Vo
FFZ
of Survey
0
44 50' N
Figure 4 Bohnenstiehl and Kleinrock: "TAG valley floor faulting"
A'
N
A'
B
B'
Backscatter Amplitude
(x 10 5 )
1.4 B.
1.2 1.0 0.8 0.6 0.4 0.2 -3870 -3875 -3880 -3885 -3890 -3895 -3900 -
b)
Heave
Throw
Bathymetric Depth
(m)
100 m
A
a)
|
50
|
100
|
150
Distance Along Profile B-B' (m)
Figure 5: Bohnenstiehl and Kleinrock "TAG valley floor faulting"
Number of Faults
18
16
14
12
10
8
6
4
2
0
East Dipping
Faults of FFZ
West Dipping
Faults of FFZ
360 5
20 35 50
Strike (degrees)
65
80
Figure 6: Bohnenstiehl and Kleinrock "TAG valley floor faulting"
Maximum Throw (m)
γ = 0.030 (TAG)
γ = 0.007(EPR)
a)
Cumulative Frequency
Fault Length (m)
D = 1.64
b)
Cumulative Frequency
Fault Length (m)
D = 1.92
c)
Fault Length + 68 (m)
Figure 7: Bohnenstiehl and Kleinrock "TAG valley floor faulting"
Throw (m)
Cumulative
Seg. D
Seg. A
Seg. C
a)
Fault 2
Distance from northern tip of fault 1 (m)
Fault 1
Segment B
Segment C
N
Accomdating fault
within relay ramp
b)
Segment A
500 m
Relay ramps
Segment D
Fault 2
TAG active
hydrothermal mound
c)
Figure 8: Bohnenstiehl and Kleinrock "TAG valley floor faulting"