Structure of the Wollaston Lake Reflector (Trans-Hudson Orogen,
Canada) from reflection AVO analysis: fractured diabase intrusion,
fluids, or silicified shear zone?
Jinfeng Ma1,2 and Igor B. Morozov2
1
State Key Laboratory of Continental Dynamics, Department of Geology, Northwest University,
China
2
Department of Geological Sciences, 114 Science Place, University of Saskatchewan,
Saskatoon, SK S7N 5E2, Canada
Address for correspondence:
Igor B. Morozov
Department of Geological Sciences
University of Saskatchewan
114 Science Place
Saskatoon, SK S7N5E2 CANADA
Tel. (306) 966 2761
Fax (306) 966 8593
E-mail igor.morozov@usask.ca
1
Abstract
A nearly 160-km long Wollaston Lake Reflector (WLR) represents one of the
most spectacular and well-recorded features imaged within the crystalline crust. It was
recorded in seismic reflection profile S2b of the 1994 Lithoprobe Trans-Hudson
Orogen transect (THOT) in northern Saskatchewan (Canada). Initial analysis of this
line targeted structural and geological interpretation and left aside the subtler properties
of reflection amplitudes. Based on modeling its normal-incidence reflectivity, the
reflector was originally interpreted as a series of subhorizontal, tabular diabase
intrusions. However, the unusual horizontal extent, nearly constant depth, compactness,
brightness, and continuity of this intrusion still remain enigmatic and suggest a
contribution of low-viscosity fluids in its formation.
To further study the properties of WLR, and also to elucidate the deep crustal
reflectors and rock properties along the Trans-Hudson Orogen transect, we reprocessed
line S2b with an emphasis on the Amplitude Variations with Offset (AVO) analysis.
While recognizing the difficulty of establishing the reflection polarity of crustal
reflections, we considered both positive- and negative-polarity interpretations. The
measurements indicate high and positive AVO gradients and suggest two possible
interpretations: 1) the reflector is caused by a massive mafic intrusion as suggested
earlier, in which case the intruded rocks should have anomalous Poisson’s ratios of
0.33, and 2) the WLR represents a silicified shear zone, with only moderate (e.g., ~
5-10%) alteration of the host rock and 0.2.
2
Although both of the above models may to some extent co-exist within the
WLR, based on its brightness, sharpness, great lateral extent and smooth shape, we
favour the second interpretation. In both models, a fractured, hydrothermal fluid zone
within a major crustal detachment could play a key role in the formation of the WLR.
The association of the reflector with laterally- and depth-migrating fluids is also
supported by magneto-telluric measurements of crustal conductivity beneath the WLR,
and also by its very long horizontal extent with only moderate depth variations.
Analogies from the studies of the Kola Superdeep Borehole (Russia) where free or
metamorphic fluids were found at comparable depths also suggest that they can play a
role in WLR formation.
Key words: AVO; bright spot; crust; Lithoprobe; Poisson’s ratio; polarity; reflection;
shear zone; Trans-Hudson Orogen; Wollaston Lake Reflector.
Introduction
As part of the Lithoprobe Trans-Hudson orogen transect, crustal seismic
reflection line S2b was acquired in northern Saskatchewan (Canada) in 1994. The line
extended for about 270 km in north-south direction, from the Athabasca basin across
the Wollaston Fold Belt and into the Precambrian rocks of the Trans-Hudson orogen
(Figure 1). One of the most spectacular crustal structures imaged by the S2b line is a
continuous band of strong reflections within the upper crystalline crust, often referred
to as the Wollaston Lake Reflector (WLR). These bright reflections extend from the
3
southern part of the complexly deformed Peter Lake domain across the Wollaston Fold
Belt and into the Middle Proterozoic Athabasca basin (Figure 2).
Several features of the WLR make it an excellent case to study the upper crustal
heterogeneities. For the subsequent discussion, we key these observations using
symbols O1-4 (Figures 2 and 3):
O1) Long, ~ 160-km long extent, with only ~ 7 km variation in depth, most
shallow under the Wollaston Lake and smoothly dipping underneath the
Wollaston Fold Belt. Similar reflectors were also observed in three
additional seismic lines ~80-100 km west of the line, (Figure 1b; I. Gyorfi,
B. Pandit, and Z. Hajnal, personal communication), suggesting that if this
structure is as continuous E_W as it is in the N-S direction, its area may
exceed ~ 17,000 km2 (dotted ellipse in Figure 1b). In addition, as Mandler
and Clowes (1998) argued, similar reflectors are found in several other areas
across western Canada.
O2) Discordance with the pervasive crustal reflectivity steeply dipping to the
north.
O3) A change in structural and reflectivity pattern and density, and apparently
in the structural style across the reflector (Figure 3).
O4) Continuity, sharpness, and brightness of its elements.
Generally, the above observations indicate a zone of some kind of structural
4
crustal contrast, probably associated with and potentially related to the formation of the
Wollaston Fold Belt. However, the nature and the source of this contrast still appear
enigmatic. Possible hypotheses suggested by the previous studies of bright crustal
reflections could be:
H1) Pockets of melts produced by recent magmatic activity (Jarchow et al.,
1993);
H2) Ductile shear zones producing thick mylonitic layers within the lower
crust (Warner, 1990) or near the brittle-ductile transition (Carmignani
et al., 1994; Nelson et al., 1996; Brogi et al., 2003), or localized
high-strain and high-porosity zones within a broad zone of ductile
flow (Ross et al., 2004);
H3) Mafic and ultramafic intrusions (Pratt et al, 1993; Vernik et al., 1994;
Mandler and Clowes, 1997, 1998);
H4) Fluid-filled porous or fractured zones (Hyndman, 1988; Warner, 1990)
which could also be associated with upper-crustal shear zones.
Of these hypotheses, H1) and H2) are unlikely in the upper-crustal, cratonic
environment, with low heat flow and abundance of seismic structure beneath the
reflector (Figure 2). Models H3) and H4), or their combination appear to be the likely
explanations of the WLR, and these hypotheses will be considered further.
In this study based on localized seismic observations, we do not expect to
5
become able to strictly differentiate between the two hypotheses H3) and H4), and
particularly because they are not mutually exclusive. Mandler and Clowes (1997), who
proposed model H3) for the WLR, explained its bright segments as caused by diabase
sills associated with the 1265-Ma Mackenzie dyke swarms that originated about 1800
km to the north of the study area, and they also pointed out outcrop evidences of
exposed Mackenzie-age sills exceeding 40 km along strike, including within the area of
Wathaman batholith (Figure 1; Hulbert et al., 1993). These observations can certainly
account for observations O4) above, yet it still remains unclear how a dyke swarm
could become confined to a narrow, ~ 50-150-m wide depth zone within >17,000 km2
of its possible emplacement, without creating a vertically more complex structure
Although the intrusions should have propagated through subhorizontal crustal fracture
zones southward, the S2b reflection section also suggests that along such a path,
numerous steeper south-and-upward leading pathways could also be present (Figures 2
and 3). Therefore, the explanation of the WLR by emplacement of mafic melts into an
initially undisturbed granitic crust appears to have difficulties from a mechanical
standpoint. The observed change in the seismic structural fabric across the WLR (O2-3)
also remains unexplained in this model.
One mechanism that could allow Mackenzie dykes to follow a strongly
confined yet extensive and horizontally contiguous path could be a major shear zone
developed by or at the time of intrusion. If such a shear zone either predated or was
coeval with Mackenzie dykes, the two models H3) and H4) could complement each
other. In terms of seismic observations, H3) could contribute the high amplitudes,
6
positive polarities, and localized character of the observed reflections while H4) better
accounts for the structural style. In this paper, we focus on hitherto unexplored
reflection amplitude variations with offset (AVO) in order to determine whether
porosity or fluids could have a significant contribution to the seismic properties of the
WLR.
Variation of reflection amplitudes with offset is among the most sensitive
indicators of low S-wave velocities, which are often associated with weakened or
alterated shear zones. Both exploration and deep crustal AVO measurements often
point to the variations of Poisson’s ratios that are often indicative of porosity, fracturing,
and presence of melts (Schmeling, 1985) and aqueous fluids (Makovsky and
Klemperer, 1999). Therefore, if observed, AVO anomalies could yield strong evidence
in favour of hypothesis H4 above.
Seismic reflectivity patterns observed in the S2b section suggest that if
hydrothermal or meteoric water is contributing to the WLR (Hyndman, 1988; Warner,
1990), such fluids would likely propagate down (and probably north) through the WLR
detachment and to some extent could penetrate into the deeper crust, possibly through
the fabric indicated by the north-dipping reflectivity (Figure 2). Note that the
reflectivity below the WLR appears somewhat stronger than above it (Figure 3),
suggesting that the WLR region could serve as a source (of fluids?) for the underlying
region but not to the region above it. Thus, a contribution of the hydrothermal or
meteoric fluids appears to be viable mechanically, provided the relation between the
7
lithostatic and hydrostatic pressures allows open cracks at the depth of WLR
(Abdrakhimov et al. 1996).
Geochemical anomalies indicated by the presence of extensive retrograde
hydrothermal mineralization due to the presence of water, saline solutions, and helium
were reported in fracture zones to the depths of about 12 km in the Kola borehole
(Kozlovsky, 1987). Seismically, such weakened, fluid-filled zones create subhorizontal
reflectivity discordant with the foliation pattern, as indicated by the studies of the Kola
Super Deep Borehole (Lizinsky and Lanev, 1991; Ganchin et al., 1998) and what could
be similar to the WLR (Figure 2). The existence of similar, foliation-controlled
microcracks to at least 4-km depth was also suggested from a petrophysical study of the
KTB borehole in Germany (Kern et al., 1991).
In this paper, we reprocess the THOT line S2b with emphasis on the AVO
analysis. In addition, we also critically re-examine the observations of reflection
polarity and consider both positive- and negative-polarity WLR interpretations. The
results show that the WLR exhibits high AVO gradients, which could be caused by
either the presence of strongly fractured, fluid-filled, high-velocity material or,
conversely, by a silicified, moderately low-velocity shear zone. For both models, on the
weight of the combined seismic evidence O1-O4) and other data, the WLR could be
best explained by a major crustal shear zone that could have provided pathways for
mineralization and emplacement of mafic intrusions, and at present, this zone should be
filled with (likely) aqueous fluids.
8
Reprocessing THOT line S2b
The initial processing of the THOT line S2b by Compagnie Générale de
Géophysique focused on structural and geological interpretations and did not consider
the amplitude properties of reflections. Because of the uncertainties in the NMO
velocity model (in our opinion, virtually unavoidable in 2D crustal profiling),
amplitude, and particularly spectral properties of the resulting stacks are strongly
affected by the quality of focusing in the NMO correction and stacking process. In
consequence, the stacked section cannot be unambiguously interpreted as representing
an idealized zero-offset (with collocated sources and receivers) recording. Instead, if
true amplitudes are of essence, careful pre-stack analysis is required.
Along with the AVO effects and geometrical spreading, numerous factors affect
the reflection amplitudes (Juhlin, 1990): shot strength, geophone coupling, lateral
heterogeneities within the overburden, absorption, tuning caused by layering, and other
effects. By contrast to using explosive sources, the use of Vibroseis sources in line S2b
resulted in a zero-phase wavelet and somewhat increased ground roll complicating the
static solutions and the subsequent analysis, yet the reflections were recorded
consistently within offset ranges corresponding to reflection angles of up to ~35º
(Figure 4). In addition to the standard reflection processing, our reprocessing included
compensation of attenuation and transmission losses, surface-consistent deconvolution,
and true amplitude recovery (Table 1). In order to obtain an accurate refraction static
solution and flatten the WLR in CMP gathers, we carefully picked the first breaks from
9
al of ~2700 shot gathers. The use of the Generalized Linear Inverse refraction statics
(GLI; Hampson and Russell, 1984; Schneider and Kuo, 1985) also helped in improving
the processing quality. Crooked-line binning size was reduced from 25×2000 m2 to
25×1500 m2 in order to reduce the effects of cross-line dips and azimuthal variations
within the CMP gathers. Four segments of the resulting stack including the WLR are
shown in Figure 2.
The AVO analysis method commonly used in oil and gas reservoir prediction is
based on a linear approximation of the Zoeppritz equations (Shuey, 1985). By weighted
stacking of the traces within the individual CMP gathers, common-reflection angle
stacks and AVO intercept and gradient sections are obtained and interpreted. However,
in deep crustal reflection records, significantly smaller portion of energy is related to
the primary reflectivity, and the AVO analysis needs to be performed in a more detailed,
per-reflector, manner (Juhlin, 1990; Makovsky and Klemperer, 1999) which should
also allow for the residual statics and focusing/defocusng effects not accounted for in
2D processing. To analyse the AVO properties of the WLR reflector, we used a simple
pre-stack approach based on the exact Zoeppritz equations (1919) applied to reflection
amplitudes normalized by the pre-reflection noise.
Approach
Using the WLR reflections picked from pre-stack CMP gathers and
disregarding the obvious amplitude outliers, we measured the AVO responses within
parts of the WLR where the data quality appears to be the best (CMPs 7340 to 7350;
10
this range was also used in normal-incidence reflectivity analysis by Mandler and
Clowes, 1997). Two distinct reflectors are present in this CMP range: at 1.7 s (~4.5-5
km depth) and 2.2 s (~6.5 km), which we labelled WLR1 and WLR2, respectively
(Figure 4). To compensate for the source, receiver, and propagation effects, we applied
reflection amplitude normalization using RMS amplitudes measured within a window
selected above the WLR1 (Figure 4).
Reflection polarity is critical for both identification of either high- or low
velocities within the reflecting zone velocity, and it also has a critical impact on the
AVO interpretation. Because the reflection polarity is difficult to establish in deep
crustal Vibroseis records, the potential effects of this uncertainty should be considered
in the interpretation. Mandler and Clowes (1997) pointed out the difficulty of
establishing the reflection polarity in the S2b line but also argued from a coincident
explosive-source THOT refraction line R3 that the WLR reflections were likely to be
positive. However, with the dominant frequencies of the reflections in line R3 of ~18
Hz, the corresponding wavelengths at WLR depth is ~340 m, which is more than
double the 70-140-m thicknesses of the WLR layers (both estimates from Mandler and
Clowes, 1997). Such layers can hardly be assumed to remain horizontal and of constant
thickness across the first Fresnel zone diameter of ~2 km (~1.6 km for the dominant
reflection frequencies of ~28 Hz in S2b line). Therefore, in the resulting multicyclic
reflected wavetrains, distinction between the top- and bottom-side reflections might no
longer be possible, and thus the notion of a definite (either positive or negative)
reflection polarity from such depth may require reconsideration.
11
In the AVO analysis below, we use the absolute (RMS) values of the amplitudes
within 100-ms windows centred on the reflectors. Therefore, the measured amplitudes
are always positive, and the predicted AVO models also need to be modified in order to
incorporate the contributions of both top- and bottom-side reflections. Assuming the
reflector zone is comparatively thin and spatially variable within the Fresnel zone, we
note that on the average, waves reflected from the top and bottom of the reflectors add
up incoherently, and the resulting reflection amplitude is:
A  
2
2
top   Abottom
 bottom  ,
Atop
(1)
where  = sin-1(pVP) is the angle of incidence, defined from the ray parameter p and
velocity VP above the corresponding boundary. By using the average elastic parameters
of the host rock and of the reflecting layer, the top- and bottom-side reflection
amplitudes (Atop() and Abottom(), respectively) can be estimated from the Zoeppritz
(1919) equations or from the corresponding matrix solutions to the plane-wave
scattering problem (Table 5.3 in Aki and Richards, 2002), Assuming a relatively small
reflectivity (and also recognising a heuristic character of this estimate), we ignore
transmission losses across the upper boundary in the calculations of Abottom().
The observed reflection amplitudes are shown in Figure 5 as functions of
reflection angles modelled within the interval velocity model derived during reflection
data processing. To highlight the AVO trends (thick solid lines in Figure 5), the
amplitudes were smoothed using a 35-points Least-Squares average filter, which also
provided estimates of the amplitude uncertainties (grey bands in Figure 5). Despite
12
significant amplitude oscillations apparently caused by focusing and defocusing
(including cross-line) effects, strong AVO anomalies are clear in the records, with
amplitudes systematically increasing with the incidence angles (Figure 5). For
comparison, the three models of diabase intrusion in granite (Table 2) predict
significantly lower AVO gradients. Even the highest- (Poisson’s ratio) model 3 (Juhlin,
1990; Table 2) lies below the AVO gradient indicated be the best-fit curves labelled
RMS in Figure 5. The two best-fit models will be discussed below, after a brief
description of the fitting criteria.
In order to improve the constraints on the Poisson’s ratios from the observed
complex AVO anomalies (Figure 5), we compared several approaches to fitting the
modelled responses to the data. A problem with the commonly used RMS misfit
measure is in its not differentiating the amplitude trends (AVO gradients) from
amplitude oscillations and random scatter. In the RMS sense, the best-fit models for
WLR1 show only ~5% variance reduction compared to model #3 (normalized 1.05
versus 1.0 for RMS, Figure 6, top; Table 2) and ~70% (1.70 versus 1.0) - compared to
the lowest-
oscillations, strong AVO gradients are still quite apparent in the data (Figure 5). A
similar limitation of RMS fitting are well known in refraction (wide-angle) seismology,
where linear time moveouts can often be identified in the background of significant
statics yet a RMS criterion would result in a large uncertainty. Alternative misfit
measures or combinations of multiple measures help achieving better constraints on the
critical model parameters (Morozova et al., 1999).
13
To define a misfit measure that would be more sensitive to the average AVO
trend compared to the scatter and localized amplitude variations, for each modelled
AVO curve, we measured its RMS misfits from the data within two ranges of the
incidence angles: 5-20º and 15-32º. For each angle range, misfit values were
normalized by their minima, resulting in normalized misfits, which we denote F1(A,)
and F2(A,), respectively. In these expressions, A is the normal-incidence reflection
amplitude,  is the Poisson’s ratio, and  denotes the set of the observed
amplitude-angle pairs. These functions attain their minima of 1 at points (A,) where
the Zoeppritz solution for reflection coefficient fits the data best within the
corresponding angular range. A similar function for the entire range of incidence angles
is shown in Figure 6 (upper plots).
By combining the two sub-range functions, a cumulative misfit measure can be
defined:
F  A,  |    F1  A,  |  F2  A,  |   .
p
(2)
where p is an auxiliary parameter controlling the sharpness of the valleys in the
resulting distribution and used mostly for plotting purposes. The appropriate choices
for p are 0.5 (corresponding to the geometric mean of F1 and F2) and 1, although a value
of p=2 could also be used to emphasize the best-fit parts of F(A,) spectrum.
Because function (2) results in a lower trade-off between the amplitude misfits
at low and high incidence angles, it is significantly more sensitive to the variations of ,
and thus it more adequately reflects the interpretative visual perception of an AVO
14
gradient (Figure 5). For example, F(A,) distributions for WLR1 with p=1 show a
140% increase in misfit from the best optimal model (dots in Figure 6) to model #3 and
220% to model #2 in Table 2.
In order to establish what levels of F(A,) could be considered significant for
distinguishing the AVO solutions, correlate the F(A,) distributions (Figure 6; p=1) to
the AVO observations in Figure 5, and compare them to the positions of labelled
models #1-3. From the observed AVO amplitude trends (Figure 5), it is apparent that
for WLR1, the best-fit model should have distinctly higher  than model #3, whereas
for WLR2, only model #2 could be considered clearly distinct from the one with F=1.
Therefore, the level of F=1.8 separating models #2 and #3 for WLR2 could be
considered as a borderline case, which also works well for WLR1. This level is
highlighted with a cross-hatched pattern in Figure 6.
Two models of the WLR
Starting from the two-interface absolute-amplitude AVO model in equation (1),
we considered two groups of models of the WLR zone: based on low-velocity (LVZ)
and high-velocity zones (HVZ). For each of these cases, by setting a reasonable
perturbation of VP, and trying a range of AVO intercepts (normal-incidence reflection
amplitudes) and values of (Poisson's ratio) within the zone, maps of normalized RMS
misfits were created, from which the best-fit points were identified (Figure 6, top). In
all cases, the exact Zoeppritz equations for reflection coefficients at oblique incidence
were used. Note that the layering of the WLR zone is likely to lead to its properties
15
anisotropic at seismic wavelengths, and hence no Poisson's ratio can be rigorously
defined for it. However we use this term in its conventional (isotropic) sense as
representing the ratio of the P- (VP) and S-wave (VS) velocities assigned to the layers of
the model:




1
1
,
  1
2


2
V

  P V  1
S 
 

(3)
This definition facilitates easy comparisons with granite (0.25) and also with most
published presentations of VP /VS data (e.g. Christensen, 1996).
The best-fit HVZ model (RMS, Figure 5) was based on the granite-diabase
model #2 (Table 2), in which the observed AVO anomalies (taken at values of
F(A,)= 1.8 for both WLR branches) required high Poisson’s ratios of the reflectors:
 > 0.37 for WLR1 and  > 0.32 for WLR2 (Figure 6, p=1). Such Poisson’s ratios are
characteristic of sedimentary environments. Note that from Figure 5, the principal
difference of the observed AVO response for WLR2 from that of WLR1 appears to be
the more pronounced amplitude oscillation (modulation) complicating the observations.
This modulation leads to a greater uncertainty region, although the Poisson’s ratio of
WLR2 is probably not significantly lower than that of WLR1 (Figure 6).
Thus, assuming a high-velocity diabase intrusion, the AVO requires values of 
exceeding ~, which are significantly higher than could be explained by competent
diabase. Such high Poisson’s ratios imply shear-wave velocities that are anomalously
16
low even for mafic lithologies of the intrusions (Christensen, 1996; Juhlin, 1990).
However, within a partially fluid-filled porous zone, shear-wave velocities could be
further reduced, resulting in high .
In an alternative (LVZ) model, by allowing a reduced VP within the WLR zone,
the observed AVO gradients could also be easily explained by increased or
near-constant VS leading to reduced  (dashed triangle in Figure 7). For crustal rocks,
both VP and Poisson’s ratio generally reduce with increased quartz content (Christensen,
1996), and therefore a model of WLR as a silicified shear zone appears to be a viable
candidate (model 4 in Figure 7 and Table 2). A broad range of LVZ model parameters
could explain the observations; for example, the AVO gradients lie nearly entirely
between those in a 10-% quartrtzite reflector model and another model differing from it
by only 0.7% in VS (Figure 5). Both of the proposed models are further considered in
Interpretation and Discussion sections below.
Note that because of the sharp differences in the two interpretations, in the
following discussions, we refer to the values of Figure 6) as the “apparent” Poisson’s
ratio. The value represents the AVO gradient modelled using the assumption of a
high-velocity (diabase; Table 2) intrusion in granite.
Interpretation
Due to inconclusive evidence for reflection polarity, two contrasting models
based on assumed positive- or negative-impedance WLR anomalies arise from the
AVO measurements above. As it is commonly done in seismic exploration industry in
17
the absence of well-log data calibration (Brown, 2005), both such interpretations
should be considered.
To summarize the AVO properties of both possible HVZ and LVZ AVO
models, Figure 7 shows the AVO fit function computed for the smoothed data for the
WLR1 reflector (Figure 5) using equation (2) with p=1. Because high density is the key
factor responsible for weak AVO responses of diabase, density variations are included
in this diagram as a quadratic function of VP and VS approximating the densities of
quartzite, granite, and diabases 1-3 in Table 2 (the function is given in Appendix A).
With density thus replaced, a continuous set of models spanning the range from
quartzite to granitic and diabase lithologies can be parameterized by VP and VS and
examined (Figure 7).
As Figure 7 shows, only the somewhat extreme HVZ model #3 (Table 2) lies
within the margin of WLR1 AVO fit. The fit could be improved if anomalously low
bulk VS (and, in fact, VP) is present (Figure 7). Reduced VP and VS could be caused by
fluid-filled fractures; for example, in the case of vertical propagation through a zone of
horizontally-aligned fracturing (Hudson, 1981; Hudson et al., 2001; Appendix B):
2

   2  
16 
~2
2
VP  VP 1 

  ,
 3(   )(1  K )  2   
16    2   
~
2
VS2  VS 1  
 .
3  3  4   

(4)
Here,  and  are the Lame parameters of the rock matrix, kf is the bulk modulus of the
fluid, c is crack half-thickness, c/a is their aspect ratio, and  is the crack density, and
18
(Hudson et al., 2001):
2
K
kf
VP
.
2
VS c (   )
a
From relations (4), the P- and S-wave velocities are affected differently by
fracturing (arrows in Figure 7). With increasing , VS generally decreases faster than VP,
and the Poisson’s ratio increases. For example, for kf = 2.2 *109 Pa and crack aspect
ratio c/a = 0.001, and  = 0.1,  reaches ~0.35 in our diabase intrusion model #2 (Table
2). The P-wave velocity becomes VP = 6.57 km/s, which still ensures a positive
normal-incidence reflection amplitude from the granite-diabase contact. Thus, a model
of fractured diabase encased in granitic wall rock could have both positive polarity and
strong AVO gradient close to the observed in the S2b data.
In addition, partial fluid saturation could lead to anti-correlated variations of VP
and VS. Although we are not aware of any similar observations for deep crustal rocks,
measurements in sedimentary rocks (Domenico, 1976; Ensley, 1984; Christensen, 1989)
suggest that with brine saturation increasing, VP could increase and VS decrease
simultaneously, leading to anomalously high Poisson’s ratios and AVO gradients.
Therefore, extensive fracturing, and particularly partial fluid saturation could cause or
enhance the observed AVO anomalies for a high-impedance WLR infill, such as a
diabase intrusion.
By comparison, the low-impedance WLR model #4 needs only a slight reduction
of VP with constant or increased VS to explain the observed AVO gradients (Figure 7).
19
In the partial-saturation model considered above, this could be achieved by the WLR
zone being somewhat drier compared to the surrounding rock.
However, this
mechanism would hardly explain the sharpness of the reflector. In another explanation,
increased SiO2 in crustal rocks is known to cause slight reduction of VP while strongly
reducing the Poisson’s ratio (dashed triangle in Figure 7). Christensen (1996) derived
an empirical relation between the Poisson’s ratio () and SiO2 content (p SiO2, in %):
pSiO2 = 100.9 – 496.92.
From this relation, an increase in silica in granite by 3% would cause reduction of the
Poisson’s ratio from 0.25 to 0.24. As an end-member model of a silicified shear zone, a
layer of quartzite (model #4 in Table 2) emplaced in granite or granitic gneiss would
lead to a strong negative reflection and AVO anomaly that is in fact somewhat too
strong for WLR1 (Figure 7).
Note that the low-impedance (LVZ) interpretation is very tolerant to its parameters
and can be easily supported with only small deviations in material properties from the
host rock. The AVO effect of pure quartzite is much stronger than the one observed, and
because the true amplitude of the reflection is unknown, a small fraction (e.g., 5-10%)
of quartzite could also be sufficient to explain the anomaly (Figure 5, top). In addition,
silicification of the shear zone is also likely to be associated with fluid-filled fracturing,
which should reduce the seismic velocities, and especially VS (arrows in Figure 7). For
example, reduction of VS by only 0.7% in the 10-% quartzite model reduces the AVO
response to the level of observed gradients (lower dashed line in Figure 5, top). Thus, a
20
broad range of LVZ models could agree with the observations (dashed triangle in
Figure 7), and a more specific set of parameters can hardly be determined.
Spatial variability
In order to characterize the spatial variability of rock properties along the
seismic line, we constructed a misfit function F() similar to the one used above for
each of the reflectors. Similarly to the above, the principal requirement to this function
is that it should be most sensitive to the local (regionally varying) values of  while
staying independent of the amplitude scaling. The following definition satisfies these
requirements:
F   
1
N
N
 F  A  ,  |   .
i 1
i
i
(2)
Here, a window of N CMP gathers is centred on the output location, and i denote the
AVO readings within each of these gathers, with i=1…N. Within each of these gathers
and for every , a misfit function F(A,) is calculated using eq. (2), and the
amplitude parameter A=Ai() is selected by minimizing this function for the chosen .
The resulting N functions are then averaged, leading to expression (2). For display
purposes, this function can be further normalized by plotting the values of F()/minF()
(Figure 8). Similarly to the misfit functions in Figure 6, this normalized function attains
a minimum value of 1 at every CMP.
Figure 8 shows the calculated function (2) at every CMP of the profile using the
AVO data from N=200 adjacent CMPs and assuming the HVZ model of the reflector.
21
The variations of the apparent Poisson’s ratio along the WLR1 reveal that the AVO
gradient is consistently positive and the Poisson’s ratios are high (0.3-0.35) within
most of its extent (Figure 8). The somewhat lower values of  for CMPs 1-3500 could
be related to their greater depths (>10 km) of the reflector in this area, leading to
narrower angular ranges in the AVO analysis and consequently less reliable estimates.
In addition, the abundance of free fluids should also drop sharply near this depth (Frost
and Bucher, 1994), and therefore we could expect lower AVO gradients in this area.
Finally, non-uniformity of the reflector, variations of seismic velocities above it,
varying recording conditions, and residual statics still cause significant scatter in these
inferred Poisson’s ratios (Figure 8).
Discussion
In the vicinity of the WLR, diabase intrusions crop out as dykes and sheet- or
sill-like bodies in the Athabasca basin and in the Wollaston fold belt (Hulbert et al.,
1993). Drill-core data from the diabase bodies in the Athabasca basin showed
individual sills ranging from 50 to 155 m in thickness (MacDougall and Maxemiuk,
1995). However, neither the S2b nor other THOT lines show indications of the WLR or
its branches reaching close to the surface. The shallowest WLR depth is about 4.5 km
near CMP 8600 (Figure 2). Stacking velocity analysis reveals no zone of high velocities
that could be associated with a magmatic source in the vicinity of the WLR. Therefore,
it still remains unclear whether the diabase outcrops found in the Athabasca basin and
in the Wollaston fold belt are related to the WLR and not, for example, to the steeper
22
dipping reflection fabric identified in the profile (Figure 2).
In characterizing the patterns of crustal reflectivity, it is important to
differentiate between bright spots and continuous semi-coherent reflectivity. Whereas
the former a likely to be associated with localized strong heterogeneities (such as
pockets of partial melts or intrusions), the latter should most likely be linked to crustal
structural contrasts, such as detachment zones and rhelogical layering. The perceived
brightness of the reflector may be relative and enhanced by its coherency and the lack
of reflectivity and scattering from the surrounding layers.
A classic example of “bright-spot” reflectivity related to tabular high-velocity
intrusions is the Siljan Ring sequence in Sweden, characterized by a series of strong,
subhorizontal to gently dipping reflections separated by non-reflective zones (Juhlin,
1990). Thicknesses of these reflectors range from a few meters to 60 m. The Bagdad
bright reflector sequence in Arizona has a remarkably similar reflection character
(Litak and Hauser, 1992). Reflection characters of the tabular mafic intrusions in both
these sequences are multi-cyclic and layered, which is probably caused by branching
during dyke emplacement. Such multi-cyclic character differs from the compact
reflection signatures of the WLR continuous across over a hundred kilometres.
In another comparison, consider the Winagami reflection sequence described
by Ross and Eaton (1997) within the Proterozoic basement beneath the sediments of
north-western Alberta. The Winagami sequence also represents an extensive suite of
high-amplitude, subhorizontal to gradually dipping reflections separated by
23
non-reflective zones at 4-19 km depths, and it was also associated with mafic or
ultramafic intrusions. The appearance of the Winagami reflections sequence is close to
the reflections in Siljan Ring and Bagdad sequences and different from the WLR.
Because the WLR is characterized by consistent reflection strength, lenticular
character, great lateral extent, and a change of structural reflectivity fabric across it, we
favour its interpretation of a crustal-scale brittle/ductile shear zone. At some time
during the orogeny, magmatic intrusions could have been emplaced into the shear zone,
similarly to the tabular intrusions and dykes found in the outcrops (Mandler and Clowes,
1997). If significant volumes of such intrusions are present within the WLR zone, they
could also be fractured and fluid-filled in order to account for the anomalously high and
positive AVO gradients. However, the seismic observations can also be explained by a
low-velocity, silicified shear zone, most likely fractured and fluid-filled, with only a
small degree of alteration and a broad range of parameters. Mylonites would also likely
have similar elastic properties. Based on the mechanically viable mechanism of its
creation (fluid infiltration and mineralization of a shear zone as opposed to tabular
magmatic intrusion), and the need for no special properties to account for the AVO
attributes, this explanation appears to be the likely case within most of the WLR extent.
Silicification is not uncommon with in the crust, and NW of the study area, it is
abindant in the sandstones of the Athabasca Basin, with quartzite ridges and
mineralization related to uranium exploration. Extensional features were inferred from
recent regional and shallow seismic investigations in this area (I. Gyorfi, personal
24
communication). Although there still is no clear evidence for relation of this
mineralization to the deep crust or for crustal-scale extension, the above model is not
excluded, if not confirmed, by this indirect evidence.
Both of the proposed WLR models could thus agree with the reflection data.
Note that in both models, crustal (likely aqueous) fluids play a major role in the
formation of the reflector. These fluids could have been driven by gravity from the
surface, released by dehydration from progressive metamorphism of the middle and
lower crustal rocks, and/or be of hydrothermal origin. However, at depths from
4.5-13.5 km, the degree of mobility of these fluids still remains unclear.
In a stable craton, gravitationally driven fluids should be able to gain access to
the upper ~ 10 km of the crust. According to Frost and Bucher (1994), within the
uppermost crust (~ 5-km depths), meteoric water could penetrate through open clefts,
usually connected to the water table. Below this level and to ~ 10 km, water could still
be present and trapped in isolated clefts, and below ~ 10 km (at temperatures above ~
250º), cratonic crust should be “dry” because of the hydration reaction rate becoming
practically instantaneous on the geologic time scale (Frost and Bucher, 1994). Thus,
petrologic and geochemical considerations do not rule out the existence of crustal fluids
at the WLR depths, at least within the southern three quarters of its extent (Figure 2). If
the reflector is at least partially associated with meteoric fluids and connected to the
water table, then it’s gently sloping shape could also be naturally explained as a less
permeable crust underneath. However, a critical point in this explanation is whether
25
sufficient bulk porosity and permeability could exist within shear zones at these depths
in order to support such fluid transport.
The primary argument against the possibility of pervasive in situ porosity and
fracturing containing free fluids at crustal depths is the expected closure of the fractures
at 7.5-9 km depths. Experimental data on shear-zone permeability at high pressures are
very scarce, and some insight could be gleaned from the observations from the Kola
Superdeep Borehole (KSDB) in Russia (Kozlovsky, 1987). KSDB penetrates a variety
of rock types that are generally subdivided into the Proterozoic (Pechenga) complex of
metasedimentary and mafic metavolcanic rocks (0-6.8 km depth) and the Archean
complex consisting of schists and gneisses (6.8-12.26 km) (Lanev et al., 1987). Starting
at the depth of 4.5 km, near the Lutchlompol fault zone, almost all of the rocks are
foliated and exhibit strong dips in the range of 20º - 70º. This depth also corresponds to
a transition from greenschist to epidote-amphibolite facies. Thus, in its major tectonic
and structural features, the setting of KSDB could be considered as a broad analog to
the THOT area.
Based on extrapolation of their measurements of the cores from KSDB at
confining pressures of up to 100 MPa, Vernik et al. (1994) argued that the observed
microcracks were entirely due to the stress relief cracking, thermal cracking, and core
disking during sample retrieval, and thus no cracks could be open in situ. The ultrasonic
velocities at these pressures were found to be close to those observed from the sonic log
and VSP data and thus they could be close to the in situ velocities. However, crack
26
closure should also depend on the relation between the lithostatic and hydrostatic
pressures, and the cracks might also re-open at depth (Abdrakhimov et al., 1996).
Similarly, because of the high pressure sensitivity of the KSDB samples, their
observed low permeabilities (Morrow et al., 1994) strongly depended on the estimates
of the fluid pressure at depth. As Borevsky et al. (1987) suggested, within parts of the
well, fluid pressure might be above hydrostatic. Fracture zones intersected by the well
to its total depth of about 12 km were reported to show geochemical anomalies,
apparently associated with the presence of water, saline solutions, and helium
(Kozlovsky, 1987).
From the above, it appears that both direct and indirect evidence from KSDB
suggest that chemically bound and even free fluids could exist at the upper-crustal
depths, particularly within the major shear zones such as the one with which the WLR
is associated in this paper. The association of the reflector with laterally migrating
fluids could also be supported by its very long horizontal extent with only moderate and
smooth depth variations, and also by the increased bright reflectivity underneath the
reflector, suggesting that propagating fluids could also penetrate from the WLR zone
into the crust underneath it. This is also supported by the recent results of
magneto-telluric (MT) imaging (Garcia and Jones, 2005). The granitic upper crust is
resistive ( > 700 Ω.m; Garcia and Jones, 2005) across the entire measured extent of the
WLR. Underneath the WLR, a layered pattern of crustal conductivity was revealed,
with conductivity increased beneath it and two conductive zones centred on the
27
north-dipping sub-WLR reflectors (Garcia and Jones, 2005). This increased
mid-crustal conductivity was associated with crustal fracturing and mineralization
(Garcia and Jones, 2005).
Interestingly, the depth to the zone of increased mid-crustal conductivity
(Garcia and Jones, 2005) correlates with the apparent Poisson’s ratio profile (Figure 8).
Poisson’s ratios are high (0.32-35) in parts of the WLR located closely to the
conductive zones within the Wollaston Domain (Figure 8). In the Peter Lake Domain,
Poisson’s ratios drop to ~ 0.30, and the depth to the conductive zone increases to ~ 4 km
beneath the reflector. Both the decreased AVO and MT responses show that the WLR
and the crust beneath it might be “drier” in its southern part. Somewhat speculatively,
this observation also suggests that meteoric fluids from the Wollaston Lake (located
east of the “wet” part of the reflector along line S2b; Figures 1 and 8) could contribute
to the increased crustal fluid content within the Wollaston Domain.
In relating the WLR to mafic and ultramafic intrusions, it is important to
correlate the S2b reflectivity to the mafic outcrops within the Peter Lake Domain
(Figure 1). South of the WLR and below Peter Lake Domain, a strong branch-shaped
reflector shallowing underneath the Wathaman Batholith was observed within the
upper-crustal section (labelled WB in Figure 2). Although its waveforms are somewhat
simpler and more continuous (compare to Figures 9 and 3), its general similarity to the
WLR was also noted (Corrigan et al., 2005; Hajnal et al., 2005). The reflector was
associated with a dyke swarm, and its apparent offset from the WLR below the Peter
28
Lake Domain (Figure 9) was interpreted as due to its relation to a different part of an
interconnected network of dykes and sills, which may be less apparent in the seismic
section beneath the WLR (Corrigan et al., 2005). In another interpretation, this reflector
was offset from the WLR by an unnamed shear zone (Hajnal et al., 2005).
Unfortunately, the strength and consistence of this reflector is insufficient for reliable
AVO analysis.
However, considering the structural character of the WB reflector, note that the
S2b line bends westward between CMPs 9000 – 13000 and crosses the Parker Lake
(PLSZ in Figure 1) and Reilly Lake shear zones (RLSZ) at oblique angles, and a
striking similarity of the reflectors with the surface traces of these shear zones can be
seen (Figure 9). Faint yet consistent north- and SE-vergent linear events can be traced
to near-zero reflection times, as indicated with conjugate arrows in Figure 9 (labelled
by the names of the corresponding shear zones and the directions of their imaging, such
as the reflection from the SE side of the Needles Falls shear zone, NFSZ-SE, Figure 9a).
These events are likely to be caused by broadside reflections from these subvertical
shear zones. Note that the NFSZ is subvertical and the RLSZ is dipping NW (Carrigan
et al., 2005), supporting the preferred imaging directions shown in Figure 9. From the
proximity of the WB reflector to the near-surface reflection from the Reilly Lake shear
zone (Figure 9b), it appears that WB could also be linked to this shear zone and may not
necessarily be related to WLR. Finally, note a very weak reflection in Figure 9c that
could come from the base of the mafic and ultramafic intrusion shown in Figure 1. If
the intrusion is dipping north, this reflection may also arrive from broadside, which
29
would account for its low apparent dip.
Information from crustal reflectors, and particularly from bright and continuous
reflectivity reflectivity within the upper crust is most important for understanding the
local structures as well as crustal genesis in general. Structures seismically similar to
the WLR were reported from several Lithoprobe projects (the Head-Smashed-In
reflector (HSI) in southern Alberta and the Winagami reflection sequence in central
Alberta; Ross and Eaton, 1997; Mandler and Clowes, 1998). Assuming that this
similarity holds (although still not tested for AVO), the conclusions drawn from the
above analysis of the WLR could apply to a combined area of at least ~140,000 km2
broadly scattered across western Canada, suggesting that a common widespread factor
controlling formation of extensive reflecting heterogeneities is likely to exist in this
large area.
Looking for such a common factor, Mandler and Clowes (1998) suggested
rheologic layering favouring emplacement of thin sheet-like mafic and ultramafic
dykes. These dykes could originate from below (with feeder dykes not observed in the
reflection seismic sections), or be emplaced horizontally and propagate to ~1000-km
distance from their source, as parts of the MacKenzie swarm (Ernst, 1995; Mandler and
Clowes, 1998). However, the extremely high Poisson ratios above suggest that even if
such a system of extremely thin dykes could be emplaced in the upper crust, it must
have also been fractured and likely altered with fluids and mineralization. The contrast
in reflectivity patterns above and below the WLR (with apparently similar contrast for
30
the HSI reflector; Mandler and Clowes, 1998) suggest that these reflectors might be
linked to a (partially) fluid-filled major brittle/ductile shear zone within the upper crust.
This shear zone could have provided the pathways for MacKenzie dykes; however, as
the analysis above suggests, the observed reflectivity could be explained by
silicification of the shear zone and does not require mafic intrusions. Therefore,
mineralization (in particular, with silica) and fluids could be another first-order
common factor defining the upper crustal structure in western Canada.
Conclusions
In most parts of the Wollaston Fold Belt and Peter Lake Domain where the
measurements could be carried out reliably, the Wollaston Lake Reflector (WLR) is
characterized by strong positive amplitude variations with offset (AVO). These
observations allow two contrasting explanations:
1) The reflector is caused by a massive mafic intrusion as suggested earlier
(Mandler and Clowes, 1997), in which case the Poisson’s ratio of the intruded
rocks should exceed 0.33. In order to account for such anomalous , the
intrusion should be pervasively fractured and fluid-filled.
2) The reflector is caused by a silicified shear zone. The observations can be
satisfied with a broad range of parameters (such as mineralization, % SiO2,
crack density and orientation, saturation, and brine salinity), and possibly with
only a moderate (e.g., ~ 5-10%) alteration of the host rock.
Although the above interpretations may to some extent co-exist within the WLR,
31
based on its brightness, sharpness, great lateral extent and smooth shape, we favour the
second interpretation. In both cases, the reflector should be linked to a (partially)
fluid-filled major brittle/ductile shear zone. The interpreted variations in the crustal
fluid content also correlate well with magneto-telluric observations (Garcia and Jones,
2005). Similar conclusions about the existence of widespread aqueous fluids and
upper-crustal depths were made from INDEPTH seismic and magneto-telluric studies
in Tibet (Makiowsky and Klemperer, 1999, Unsworth et al., 2004, 2005). Subject to
their checking by AVO analysis, the above conclusions from WLR could potentially
relate to other similar reflectors found in several areas within western Canada.
Analogies from the studies of the Kola Superdeep Borehole penetrating the Baltic
Shield in northern Russia also suggest that free or metamorphic fluids could be present
at these crustal levels. In an additional conclusion, we note that structural interpretation
of the WLR and reflectivity along the Trans-Hudson and potentially other crustal
profiles could be improved by incorporating out-of plane reflections from the major
shear zones and structural contrasts mapped on the surface.
Acknowledgements
J.M. wishes to express his gratitude to the Deep Seismic Sounding Group at the
University of Saskatchewan. Thanks to Prof. Zoltan Hajnal, Dr. Bhaskar Pandit, and
Steven Gyorfi (University of Saskatchewan) for their advice and numerous discussions.
Special thanks to Prof. Scott Smithson (University of Wyoming, USA) for his
encouragement and information about the Kola borehole studies. The authors also
32
thank the referees, and particularly Prof. Ron M. Clowes for numerous comments on
the manuscript. This research was supported by the Chinese Scholarship Council,
Chinese National Science Foundation Grant H40144017, 40474045, 40674041, and in
part by Canada NSERC Discovery Grant RGPIN261610-03.
33
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41
Tables
Table 1. Main steps of reprocessing the THOT line S2b
1
2
3
4
Operation
Crooked-line geometry
Air-blast noise attenuation
True-amplitude recovery
Deconvolution
5
Elevation statics
6
Refraction statics
7
8
9
10
11
Surface-consistent deconvolution
Residual statics
Velocity analysis
NMO correction
AVO analysis
12
DMO, stack
13
Time migration
Summary of parameters
25×1500 m2 bins
2400 ms window, 330 m/s
1/(t*V2) scalar and 0.5 db/s compensation
Minimum-phase predictive, 60-ms lag,
4-5-80-90 Hz pass band
NMO datum, 6000 m/s replacement
velocity
GLI (Hampson-Russell); surface-consistent
model
0.5-8 s window, ±10 ms range
Iterative with NMO/DMO
30% stretch mute
Angle-limited CMP gathers at 0-45º;
maximum offset 12000 m,
Kirchhoff DMO algorithm; Power stack
normalization
Kirchhoff algorithm
42
Table 2. Three models of the WLR intrusion, listed in the order of increasing Poisson’s
ratio of diabases (), country rock, and a 100%-quartzite model
#
Country rock
(granite)
Parameters
Source
VP=6.18 km/s,
Christensen and Mooney (1995)
Vs =3.57 km/s,
=0.25,
=2.65 g/cm3
Diabase
intrusion
1
VP =6.71 km/s,
Christensen (1996)
Vs =3.73 km/s,
=0.277,
=2.94 g/cm3
2
VP =6.65 km/s,
Mandler and Clowes (1997)
Vs =3.55 km/s,
=0.3,
=2.95 g/cm3
3
VP =6.50 km/s,
Juhlin (1990)
Vs =3.25 km/s,
=0.333
=3.00 g/cm3
Silicified
shear zone
(Quartzite)
4
VP =6.01 km/s,
Christensen (1996)
Vs =4.04 km/s,
=0.085,
=2.65 g/cm3
43
Figures
Figure 1. a) Map of the Lithoprobe Trans-Hudson Orogen Transect line S2b in NE
Saskatchewan. Main tectonic units and reference common midpoint (CMP)
locations are labelled. Reference CMP locations are indicated by black circles
and labelled. B) Location mop of the profile in western Canada. Alberta (AB),
Saskatchewan (SK), Manitoba, and the Northern Territories (NT) are labelled.
Crosses are the positions of high-resolution seismic lines in the Athabasca Basin
also showing a consistent bright reflector at crustal depths.
Figure 2. Parts of the reprocessed S2b stacked section, with the Wollaston Lake
Reflector (WLR) indicated. The panels overlap by 100 CMPs. In the second
bottom panel, two branches of the reflector used in the discussion are labelled
WLR1 and WLR2. CMPs are labelled as in Figure 1. Note the strong and
coherent reflectivity and continuity of the WLR and also the Wathaman Batholith
(WB) reflector that is also thought to be related to WLR (Corrigan et al., 2005;
Hajnal et al., 2005).
Figure 3. Zoom-in into a segment of the section in Figure 2 between CMP 7200-8200.
Note the difference in the reflectivity patterns above and below the WLR,
44
characterized by: 1) less coherent pattern with steep linear reflectors (A) above
the WLR, 2) more coherent, lower-frequency, arcuate features (B1) or
lower-angle reflectors (B2) below the reflector.
Figure 4. NMO-corrected CMP gather 7340 included in AVO analysis (Figures 5 and
6). The traces were normalized using the amplitude normalization window (grey
dashed lines). Note that the amplitudes of both reflections consistently increase
and oscillate with the offsets.
Figure 5. AVO responses of WLR1 and WLR2 from CMPs 7340-7350. Small crosses
are the measured amplitudes; the thick solid line represents the averaged
amplitude, and the grey band – its uncertainty. Numbered lines show AVO
responses for the corresponding models in Table 2, and the lines labelled RMS are
the best-fit curves in a least squares sense, derived by using the HVZ model 2
(Table 2) and varying the Poisson’s ratio (Figure 6). Note that the AVO trends
require higher values of , and even the RMS curve clearly underestimates the
general AVO trend. Dashed lines in the upper plot represent the AVO responses
of LVZ models: the 10-% quartzite model (labelled 4) and another model with
S-wave velocities reduced by 0.7%. Note the broad range of LVZ models that can
fit the AVO data.
45
Figure 6. RMS (top) and the modified misfit functions F(A,) (eq. 2) for p= 0.5
(middle) and p= 1 (bottom). Labelled dashed lines correspond to the three
intrusion models in Table 2, and black dots indicate the RMS and F(A,) best-fit
solutions. The cross-hatched pattern within the contour F(A,)=1.8 in the bottom
plots indicate the interpreted model uncertainties. Note that the F(A,) criterion
provides a significantly improved discrimination between the solutions, and
plotting with increasing p emphasizes the best-fit models.
Figure 7. Two-sided reflection AVO fit for a range of HVZ and LVZ models computed
using expression (1) for smoothed WLR1 data (Figure 5). Labelled circles
correspond to four models in Table 2. Density values were approximated as
described in Appendix A. Arrows represent variations in the seismic velocities
due to 0- to 5-% volumes of brine-saturated fractures (black and grey arrows
correspond to S-wave velocities across and parallel to the direction of fracturing,
respectively; the estimation method is described in Appendix B). Dashed triangle
indicates an area of partially-silicified, fractured intrusion that appears to be in the
best agreement with AVO data.
Figure 8. Variations of the apparent Poisson’s ratio (assuming the HVZ model) of
WLR1. Note the high values of  0.33 within the Wollaston Domain decreasing
within the Peter Lake Domain and into the Athabasca Basin.
46
Figure 9. Upper-crustal segments of the stacked section of S2b: a) in the vicinity of the
Needles Falls shear zone, b) crossing the Reilly Lake shear zone, c) crossing the
ultramafic intrusion within the Peter Lake domain (Figure 1). Pairs of arrows
indicate likely broadside reflections from the shear zones and from the base of the
intrusion. The WLR reflector is also indicated.
47
Appendix A: Density replacement function used in AVO
modeling (Figure 7)
In order to reduce the number of independent variables to two and interpolate
between the quartzite, granite, and diabase models (Table 2) in Figure 7, we
approximated the density using the following function:
  2.65  0.68 * (4.0 - VS ) * (VP - 6.18) - 0.39 * (4.0 - VS ) * (VS - 3.57)
in which the density is measured in g/cm3 and seismic velocities – in km/s. This
formula provides an approximation suitable for illustration of a range of LVZ and HVZ
models fitting the AVO data (Figure 7).
Appendix B: Effects of crack porosity on bulk seismic
velocities (eqs. 4)
To estimate the dependence of velocities on crack volume (eqs. 4), we use the
model by Hudson (1981) and Hudson et al. (2001). Their effective-medium equation
for the P-wave velocity in fluid-filled rock VP reads:
~
V p2
VP
2
 1  P(
2
2
2


VP 2
VP
PVP
4
)


P
(
1

)
sin


(
Q

) sin 2  cos 2   ,

2
2
2
2VS
VS
VS


where VP and VS are the wave speeds in the isotropic matrix material,  is the angle of
wave propagation relative to the crack planes, and:
48
16 
(1  K ) 1 ,
3(   )
64 
Q
,
3(3  4  )
P
2
K
kf
VP
.
2
VS c (   )
a
Here,  and  are the Lame parameters for the matrix material, kf is the bulk modulus of
the fluid, c is crack half-thickness, c/a is their aspect ratio, and  is crack density
(Hudson et al., 2001).
For S waves polarized across and along the crack plane, the velocities are
respectively (modified after Hudson et al, 2001):
~
 (1  K ) 1
Vs  2
16
  2
16
4  2
2
( ) 1  (
)   (  2  )  

 sin  cos  ,
VS
3 3  4 
3
3  4  
 
~
Vs|| 2
16
  2
16
  2
( ) 1  (
) (
) sin 2  .
VS
3 3  4 
3 3  4 
By setting =0, eqs. (4) in the text are obtained.
49