Intraplate Earthquakes and Postglacial Rebound in Eastern Canada

Intraplate Earthquakes and Postglacial Rebound
in Eastern Canada and Northern Europe
Patrick Wu
Dept. of Geology & Geophysics, University of Calgary,
Calgary, Alberta T2N-1N4, Canada,
[email protected]
Keywords: Earthquakes, Faulting, Mantle viscosity, Seismotectonics, Stress distribution
Abstract. The causal relationship between postglacial rebound and intraplate earthquakes in
Eastern Canada and Northern Europe is investigated with the finite element model. Prominent
features of this analysis are the inclusion of: i) a stratified, viscoelastic mantle, ii) a realistic
deglaciation model and iii) the ambient tectonic stress and overburden stress contributions in
the calculation of the total stress field. It is demonstrated that the spatial distribution of current
seismicity in these areas cannot be explained by the strain rate distribution due to rebound. In
order to explain the observed spatial distribution of earthquakes, the mode of failure and the
timing of the pulse of earthquake/faulting activity following deglaciation, both postglacial rebound and tectonic stress are required. In addition, the observed orientations of the contemporary stress field and the rotation of stress in E. Canada since deglacial times can be explained by
a viscoelastic Earth with uniform 1021 Pa-s mantle. The effect of a high viscosity lower mantle
has also been investigated. It is demonstrated that a high viscosity lower mantle will introduce
zones in E. Canada where earthquake activities will increase in the future. Also any significant
rotation in stress orientations since the last deglaciation is prevented. Otherwise, it has no effect
on the mode of failure and does not significantly affect the onset time of intraplate earthquakes.
1. Introduction
Large intraplate earthquakes are found in Eastern Canada (Laurentia) and Northern Europe
(Fennoscandia). In order to mitigate more effectively the hazards associated with these earthquakes
and to plan for safe storage of nuclear toxic-waste in underground repositories, it is vital to understand
the spatio-temporal variation of the state of stress, the fault potential and the cause of such earthquakes.
A fundamental question in the study of intraplate earthquakes in Laurentia and Fennoscandia
is the relative importance of plate tectonics and postglacial rebound (including both glacial loading
and unloading) in earthquake generation. There are geological and geophysical evidence that support
postglacial rebound as the dominant cause of these intraplate earthquakes but there are other evidence
that favors tectonic stress as the dominant cause. These evidence will be reviewed in the next section.
To resolve the issue, Wu & Hasegawa [51, 52] and Wu [49, 50] have used the finite element method to
model the spatio-temporal distribution of stress and changes in fault stability in E. Canada. The aim of
this chapter is to review their work and extend the analysis to N. Europe.
Prominent features of the present series of studies [49, 50, 51, 52], which distinguish them
from previous ones [20, 35, 40, 41, 48] are the inclusion of : (1) a viscoelastic mantle and thus the
migration of stress associated with viscoelastic relaxation; (2) a realistic ice sheet deglaciation history
and sawtooth cycles of loading and unloading; (3) glacial/deglacial induced stress, tectonic stress and
overburden pressure contributions in the calculation of the total stress field; and (4) deviatoric stress
and mean stress in the computation of fault instability [21, 22].
The plan of the chapter is as follows: After a review of the observations and the evidence for
or against rebound stress as the trigger to intraplate earthquakes, we will describe the model more
closely. Then we shall show that with a uniform viscosity mantle, most of the observations can be
predicted. Finally, we shall explore the effect of a high viscosity lower mantle on the predictions of
the model before giving the conclusion.
Earthquake location and magnitude
B
BB
af
fin
Is
la
BU
nd
LS
BA
Hudson Bay
d
Shiel
Canadia
n
OB
G
V
SL
GB
M>6
(1663-1990)
6 > M > 4.5
(1980-1990)
4.5 > M > 3.0
(1980-1990)
ST
Seismotectonic
trends
Figure 1a. Spatial distribution of the larger earthquakes and the seismotectonic trends in E. Canada.
Symbols are as follows: BU- Bothia Uplift; BA- Bell Arch; BB - Baffin Bay; LS-Labrador
Sea; GB - Grand Banks; SLV- St. Lawrence Valley; OBG - OttawaBonnechere Graben.
Reprinted with permission from [52] © Blackwell Science Ltd.
II. Geological and Geophysical Evidence
Here we wish to review the geological and geophysical evidence on intraplate earthquakes in
Laurentia and Fennoscandia. Some of these observables will be calculated and compared with the
observations which act as constraints for our model.
There are two pieces of observational evidence that suggest plate tectonics may be the dominant cause of these earthquakes: First of all, the spatial distribution of recent earthquakes in E. Canada
and N. Europe shows little correlation with the center of postglacial rebound - instead, seismic activity
mainly lies along three pre-weakened tectonic zones (Fig. 1a) in Eastern Canada [4, 52] while in N.
Europe most of the larger earthquakes (magnitude > 4) are distributed along the coastal regions (Fig.
1b) while the interior is relatively nonseismic with most magnitudes less than 4 [10, 38, 39, 47].
Figure 1b. Spatial distribution of earthquakes in Fennoscandia from 1965-1995. ( Courtesy of
Institute of Seismology, University of Helsinki)
The second piece of evidence is from the contemporary stress field whose orientation does not
appear to be dominated by the effects of past glaciation - in fact, for most parts of continental E.
Canada (east of the Cordillera), the maximum horizontal stress component SHmax is almost uniformly
aligned along an ENE-NE azimuth (Fig.2a) [1, 5, 55, 56] and can be most readily explained by stresses
associated with spreading at the mid-Atlantic Ridge [36]. In Fennoscandia, the first order SHmax
orientation below 300 m depth is along NW-SE, in agreement with the direction of ridge-push from
the Mid-Atlantic (Fig. 2b), while that in Norway is affected by the Caledonides [11, 16, 42, 43]. Thus,
rebound stress has little influence on contemporary stress orientation - again indicating that the dominant stress is tectonic.
On the other hand, the good correlation between the onset time of postglacial faults/earthquakes in E. Canada and Fennoscandia with the end of deglaciation, indicates that postglacial rebound
may have played a much more important role in earthquake generation at early postglacial time. For
example, the small thrust faults in southeastern Canada offset scour marks left by the glaciers indicating that their time of formation is postglacial. The dating of some earthquake-triggered mud
slumps in S.E. Canada [37] further indicates that earthquake activity was generated at the end of
40
?
Ungava
Pen.
bra
do
r
NW Newfoundland
AT
GL
AC
IA
L
Charlevoix
MA
XI
MU
M
Gre
at
Lak
es
Indian
60W
40N
La
Cartwright
100W
12
N
MARGIN
LAURENTIDE ICE
0W
N
50
W
60
Me lvi lle
Pen.
SHmax Orientation at
early Postglacial time
Current Orientation
a
Figure 2a. Map of Eastern Canada showing: (i) orientation of the contemporary regional stress field
(bold inward-pointing arrows) on land and off-shore Labrador. Question mark near the
arrows indicates that the orientation is uncertain since only a few observations are available; (ii) orientations of an earlier stress field duduced from postglacial faults (thin inward
pointing arrows); (iii) the location of the ice margin at the last glacial maximum and (iv)
the location of some sites in Fig.7 where the evolution of of stress and dFSM are shown.
deglaciation around 9,000 years ago. The dating of earthquake-induced liquefaction features in the
Wabash Valley bordering Indiana and Illinois (USA) also gives a postglacial time of 8 to 1 ka BP [30].
In N. Europe, postglacial faults (Fig.2) were also formed between late glacial to early postglacial
times 8,000 to 9,000 years ago [6, 26] although their displacements are much larger - averaging about
15 meters and they may have fractured the whole crust [6].
All of the postglacial faults in E. Canada and N. Europe are thrust faults and are consistent
with the fault mechanism predicted by postglacial rebound. Further suggestion of a possible link to
postglacial rebound is from the orientations of the postglacial thrust faults in southeastern Canada [3]
which indicate that the maximum horizontal principal stress (SHmax) orientation of the paleo-stress
field was mainly in the NW-SE direction (Fig.2a) - perpendicular to the ice margin at glacial maximum and consistent with the direction of ice retreat [1]. Note that this orientation of the paleostress
field is almost perpendicular to the SHmax orientation of the contemporary stress field indicating that
large stress rotation occurred during postglacial time. In Fennoscandia, the postglacial faults are generally NNE-SSW trending [6, 24, 25, 27, 28] and are also subparallel to the ice margin at glacial
maximum. However the orientation of the paleostress field implied by these postglacial faults is also
subparallel to the tectonic stress caused by spreading of the North Atlantic Ridge. Thus, it is not clear
20E
0E
10
W
40E
70N
N
Vesteralen
Ga
a
lliv
anlan
d
0E
rm
Ange
re
s
Hel
Oslo
ink
i
60N
10E
40E
20E
Orebro
55N
30E
65
Figure 2b. Map of Fennoscandia showing (i) locations and orientations of some postglacial faults.
Paleostress orientations are approximately perpendicular to these postglacial faults; (ii)
orientations of the contemporary stress field; (iii) location of some sites in Fig. 7.
if paleostress orientation in N. Europe was determined by postglacial rebound or by plate tectonics.
Furthermore, Ekman [12,13] found that there is a high spatial correlation between
microseismicity, boulder caves and the maximum curvature of uplift. Basham et al. [8] and Hasegawa
& Basham [17] also found that the steep gradients in postglacial rebound contours in Baffin Island
correlate well with bands of intense seismic activity. However, earthquake clusters are not found in all
places with steep gradients in postglacial uplift.
Finally, the focal mechanisms of the larger earthquakes in E. Canada today are dominantly of
the thrust type and are consistent with that predicted by postglacial rebound. However, there are some
exceptions along the northeast coast of Baffin Island, where focal mechanisms are of the normal-fault
type, and in the northeastern United States, where strike-slip faulting [18, 19, 44, 55] are interspersed
with thrust faulting [19, 29]. In N. Europe, some of the larger recent events show thrust faulting [47],
however, the mode of failure of the majority of smaller earthquakes are a mixture of strike-slip, thrust
and normal faulting [6, 39, 47], with strike-slip motion being consistent with the notion that tectonic
stress is dominant.
To resolve the issue of which is the dominant cause of intraplate earthquakes in E. Canada and
N. Europe, it is proposed here that both tectonic and rebound stress are required: past tectonic processes are responsible for creating the pre-weakened tectonic zones and the ridge-push forces at the
Mid-Atlantic bring the faults close to failure. On the other hand, the stress induced by glacial unloading, although not large enough to dictate the location of new faults [35], can reactivate those preexisting faults that have optimal orientations. However, the stress due to glacial unloading gradually
diminish with time, so that at present, tectonic stress difference is large enough to dominate the stress
orientations. However, this does not mean that rebound stress cannot trigger intraplate earthquakes
today. In fact, rebound stress is still responsible for most of the seismicity in E. Canada and some of
the activities in N. Europe today.
III. The Model
Due to the loading and unloading of the earth (by ice sheets and melted ice-water loads) and
the relaxation of stress associated with the creep of mantle rocks, the state of stress inside the earth
constantly changes even if it is assumed that tectonic stress and all other stresses remain constant.
(This latter assumption is valid for time scales of a few glacial cycles.) As the state of stress changes,
the Mohr circle moves closer to and/or away from the line of failure according to a time-dependent
quantity called dFSM [Wu & Hasegawa 1996] which is defined by:
dFSM (t ) =
1
2
σ 1(t o) – σ 3(t o) – σ 1(t ) – σ 3(t )
+ µ β σ 1(t ) + σ 3(t ) – σ 1(t o) + σ 3(t o)
(1)
where
β = sin arctan µ / 2 µ
(2)
and µ is the coefficient of friction taken to be 0.6; t is the time under consideration; t0 is the time before
the onset of glaciation; σ1, σ2 and σ3 are the maximum, intermediate and minimum (compressive)
principal stress respectively. The physical meaning of dFSM is the change in time of the Fault Stability Margin (FSM) [21, 22], which is the shortest distance between the Mohr circle and the line of
failure (see Fig.3a). Equation (1) therefore states that FSM will decrease if the increase in deviatoric
stress, which increases the radius of the Mohr circle and makes the Mohr circle closer to failure (e.g.
Fig.3c), is able to dominate over the increase in mean stress, which moves the center of the Mohr
circle away from the line of failure (e.g. Fig.3b). Therefore, if the pre-existing faults are initially close
to but not at failure, then, a negative value of dFSM would mean enhanced likelihood of faulting
whereas a positive value of dFSM indicates that fault stability is enhanced.
The Mode of Failure depends on which of the principal stresses is closest to the vertical (see
Fig.3a). If σ1 is nearly vertical, then the mode of failure is normal; if σ3 is close to the vertical, then
a ) Fault Stability Margin & Mode of Faulting:
τ
τ=
τ0
FSM
σn
+µ
τ0
r = (σ1− σ3 ) / 2
2θ
σ2 (σ1+ σ3 ) / 2
Fault Type: σ3
Normal:
horiz horiz
horiz
Thrust:
vert
Strike-Slip: horiz v e r t
σ1
σn
vert
horiz
horiz
b) Fault Stability Margin increases at
glacial maximum
τ
τ=
µ
τ 0+
σn
FSM
'
τ0
σ3
σ1
σ3'
σ1'
σn
c) Fault Stability Margin decreases after
deglaciation
τ
τ=
τ0
σn
+µ
FSM
'
τ0
σ3 σ'3
σ1
σ1'
σn
Figure 3. a) Mohr circle, line of Failure, definition of Fault Stability Margin and the mode of failure.
b) During glacial loading,the increase in mean compressive stress moves the Mohr circle
away from failure. c)After removal of the ice, the minimum (vertical) principal stress return
to its initial value, but the maximum (horizontal) principal stress relaxes with time. The
increase in deviatoric stress and thus the radius of the Mohr circle, moves it closer to failure.
Reprinted with permission from [52] © Blackwell Science Ltd.
thrust faulting occurs; On the other hand, if σ2 is nearly vertical, then strike-slip motion occurs .
For simplicity, the total stress is assumed to be composed of the rebound stress (due to loading
and unloading of the ice sheets), tectonic stress and overburden pressure - the last two can be taken as
time-invariant in the time period under consideration. Although fluctuation of pore fluids do affect
fault stability, their time dependence is more difficult to model and thus will be neglected in this
preliminary study.
The orientation of the tectonic maximum horizontal principal stress is assumed to be determined by ridge-push at the mid-Atlantic Ridge. Thus, it is taken to be along the N60oE direction in
North America [3, 5, 55, 56], and along the N60oW direction in Northern Europe [11, 42, 43]. Although the magnitude of the maximum and minimum horizontal tectonic stresses are largely unknown, Wu & Hasegawa [51, 52] have shown that tectonic stress magnitudes have little effect on
dFSM. Furthermore, stress orientation only depends on horizontal stress differences alone [50]. In this
chapter, the maximum horizontal tectonic stress is taken to be 150 MPa - this is much larger than the
50 MPa assumed in Wu [50] and Wu & Hasegawa [52] and give stress levels closer to that observed
in Europe [54]. For the minimum horizontal tectonic stress, a range of values will be considered.
Rebound stress is calculated with the finite element method. The earth model considered here
is a compressible, stratified flat earth that consists of an elastic lithosphere on top of isotropic, viscoelastic
ρ (kg/m^3), VP (m/s), VS (m/s)
15000
VP
10000
VS
ρ
5000
A)
Elastic Structure
of Earth Models
0
0
1000
2000
3000
24
Log10(Viscosity)
B) Viscosity Models
23
L2
22
L1,L2
21
L1
20
0
1000
2000
3000
Depth (km)
Figure 4. Elastic structure and viscosity profiles of the stratified earth models L1, L2 & L3. The
lithosphere is 100 km thick under N. America but 80 km thick under N. Europe.
Maxwell layers in the mantle that, in turn, overlie an inviscid fluid core. The elastic structure and
viscosity profiles of the earth models considered are shown in Fig.4. Model L1 has a uniform 1x1021
Pa-s mantle. Models L2 is different from L1 only in the lower mantle (below 670 km depth), where
the viscosity is 1x1022 Pa-s. For both earth models, the lithosphere is assumed to be 100 km thick
beneath N. America, but 80 km beneath N. Europe. However, it can be demonstrated [50] that the
conclusions of this paper is not affected by the thickness of the lithosphere.
The initial state of the Earth is assumed to be deglaciated and without any bending stresses only tectonic stress and overburden pressure exist in the initial state. Here, loading and unloading of
the Laurentian, Cordillera, Innuition, Greenland and Fennoscandian Ice Sheets plus eustatic loading/
unloading in the ocean floor are included (see [52] for details). The ice model is adopted from the
ICE3G model of Tushingham & Peltier [45, 46]. The adopted ice thickness variation in North America
and Greenland is given in an earlier chapter by Wu & Johnston [53], and its thickness in Fennoscandia
is contoured in Fig.5. Thirty Glacial cycles are included before the final deglaciation which began
around 19,000 BP. For simplicity, the glacial cycle are assumed to have a saw-tooth shape, with slow
600
200
1000
900
1800
1200
300
1400
2100
1400
2100
1500
200
300
t = 18 KaBP
t = 12 KaBP
200
800
150
1000
50
800
200
200
400
t = 10 KaBP
t = 8 KaBP
Figure 5. Adopted version of ICE3G deglaciation history in Northern Europe . Thickness of ice is
contoured in meters.
buildup time of 90,000 years but a rapid deglacial time of 10,000 years.
IV. Results
In the following, all results of fault stability and related calculations are shown at a representative
depth of 12.5 km depth. However, all stress orientation and stress rotation calculations are for the
Earth’s surface - otherwise the large overburden pressure will make SHmax indistinguishable from
Shmin. It can be shown that the conclusions of this paper are not affected by depth in the top 40 km
[51].
IV.a Model L1 - uniform viscosity mantle
First consider the earth model that has a uniform 1x1021 Pa-s mantle. Figure 6 shows the
spatio-temporal evolution of dFSM in Laurentia and Fennoscandia predicted at a depth of 12.5 km.
Inspection of Fig.6 shows that maximum fault stability of 14 MPa in Laurentia and 8 MPa in
Fennoscandia are promoted underneath the load and in the surrounding areas at glacial maximum
around 18 ka BP (thousand years before present). Fault stability is promoted inside the ice margin
because the increase in the mean stress due to the weight of the ice has a more dominant effect than
that due to the increase in the deviatoric stress (see Equation 1).
At 9 ka BP, stability is still promoted underneath the existing ice and around the peripheral
bulge in Laurentia while instability is promoted in deglaciated areas. Instability is promoted in
deglaciated areas because the horizontal flexural (rebound) stress, which adds to the maximum principal stress, remains large at 9 ka BP. However, the minimum principal stress, which is nearly vertical,
returns to value in the initial state since the ice load is now removed. Thus the increase in radius of the
Mohr circle (relative to the initial state) has a more dominant effect than the increase in mean stress thus moving the Mohr circle closer to the line of failure. On the other hand, stability is promoted in
areas around the peripheral bulge despite of the absence of ice loads there. This is because rebound
stress, which is tensional at 9 ka BP around the bulge, causes the (compressive) maximum principal
stress to decrease while the minimum principal stress, which is nearly vertical, remains the same as
the initial state. Thus, the radius of the Mohr circle decreases (relative to the initial state), and fault
stability increases.
The situation is different in N. Europe: at 9 ka BP, when about 300-500 meters of ice still
covered the center of Fennoscandia, instability of about 2 MPa is promoted underneath the ice, while
stability is promoted in the surrounding peripheral bulge. This early promotion of instability before
complete deglaciation is due to the amplification of stress as the wavelength of the load approaches
the thickness of the lithosphere [23] and the characteristic length of the Fennoscandia load is closer to
the thickness of the lithosphere than the one in Laurentia.
For the present (0 ka BP), fault instability of about 1 MPa is promoted in the center of rebound
in Eastern Canada and northern Europe. This magnitude of fault instability is probably too low to
cause fracture, but is large enough to reactivate optimally oriented pre-existing faults. (It is larger than
the stress level that has triggered the Landers earthquake [15]). Therefore, to explain the observed
spatial distribution of earthquakes (Fig. 1), one needs to assume that the initial Fault Stability Margin
is everywhere greater than 1 MPa, except in the pre-weakened tectonic zones of E. Canada and the
coastal areas in Fennoscandia, where the initial value of FSM is close to zero.
dFSM at 18 ka BP
dFSM at 18 ka BP
3
6
8
6
2
2
7
2
2
N
7
0
6
4
8
6
5
4
14
5
3
6
12
4
10
6
0W
3
8
10
12
2
1
4
2
10
0
0
8
8
60
N
10
4
2
7
0
6
12
14
5
12
14
4
3
8
10
4
2
6
8
2
6
N
40
4
W
2
(a)
dFSM at 9 ka BP
4
3
-1
4
6
1
4
2
1
- 2
2
1
3
1
1
2
0W
- 1
- 2
0
2
4
0
0
1
2
2
N
-1
0
N
2
0
2
3
1
(d)
9 ka BP
1 2
60
1
6
dFSM at
N
60
40
0
10
12
2
- 1
- 2
1
1
1
1
- 3
12
- 2
- 1
0
-2
40
0
0
0 ka BP
dFSM at 0 ka BP
1
-1
0
2
4
6
0
4
0
2
1
60
-1
N
0
(e)
-1
N
dFSM at
0
W
(b)
1
N
0
0.0
0.5
0
N
-1
0
60
40
1
0
-1
6
0
-0.5
0.0
0
0
-1
-1
0.0
-0.5
12
0W
-1.0
-1.0
0
-1
-0.5
-1
0.0
-1
-1
40
N
W
(c)
N
0.0
60
40
0.0
(f)
Figure 6. Spatio-temporal variation of dFSM in N. America and Northern Europe at three time periods at a seismogenic depth of 12.5 km depth. Model L1 is used. Contours are in MPa.
The evolution of dFSM, the total stress (expressed in terms of horizontal principal stresses
Hmax, Hmin and the vertical principal stress PrinZ) and the mode of failure at several sites in North
America and Northern Europe are shown in Fig.7. Inspection of Fig.7a shows that for the sites along
the Boothia Uplift-Bell Arch (Melville Pen. and Ungava Pen.) and the St. Lawrence Valley-Ottawa
Bonnechere Graben (Charlevoix), dFSM becomes negative around 7-9 ka BP and reaches a minimum
value around 7-4 ka BP. The mode of failure is predicted to be thrusting since the least principal stress
is nearly vertical. This is consistent with the observation that postglacial faults in southeastern Canada
are of the thrust type [2]. In Newfoundland, where melting is early, instability is predicted to have
begun much earlier at around 13 ka BP - however, due to the loading of the nearby ocean floor by a
600
550
500
Melville
600
Hmax
Hmax
Hmin
550
Hmin
Pen.
500
Ungava Pen.
450
450
PrinZ
PrinZ
400
400
1 0
5
0
- 5
dFSM
dFSM
Thrust
1 0
5
Thrust
0
- 5
600
dFSM &
Stress (MPa)
600
Hmax
Hmin
550
500
Hmax
Hmin
500
Labrador Sea
near Cartwright
NW NewFoundland
450
PrinZ
400
2
1
0
- 1
dFSM
1
0
- 1
- 2
Thrust
Thrust
dFSM
600
Hmax
Hmax
Hmin
550
Hmin
550
500
Charlevoix, Que.
450
400
1 0
5
0
- 5
450
PrinZ
400
600
500
550
Indiana
450
PrinZ
PrinZ
400
dFSM
- 1 5
dFSM
Thrust
- 1 0
- 5
Time (ka BP)
5
Thrust
0
0
- 1 5
- 1 0
- 5
0
Time (ka BP)
Figure 7a. Evolution of the horizontal principal stress (Hmax, Hmin), the vertical principal stress
(PrinZ), dFSM and the mode of failure for six sites in Fig.2a - all at a seismogenic depth
of 12.5 km. The Earth model L1 is used in this calculation.
rapid influx of melted ice water, instability is suppressed temporarily from 10-7 ka BP. The loading of
the nearby ocean floor also causes delay in the onset of instability along the Labrador coast (e.g.
Cartwright). In areas beyond the ice margin in the northern United States, seismic and faulting activity is predicted to have occurred later at around 8 ka BP. The predicted onset time of instability in
Indiana also coincides with the observed time of a very large (MW7.5) Wabash Valley earthquake
discovered and dated by paleo-liquefaction research [30]. The mode of failure is again predicted to be
thrust faulting. Examples of pure thrust-fault earthquakes are the MW5.4 southern Illinois earthquake
of 1968 [19] and the mb5.1 Goodnow earthquake in New York State [29].
For Fennoscandia, the early promotion of instability due to stress amplification results in an
550
550
500
450
400
500
Hmin
Hmin
Oslo,
Norway
450
Orebro,
S. Sweden
400
PrinZ
PrinZ
350
300
1 0
5
0
- 5
Hmax
Hmax
dFSM
dFSM
350
300
1 0
5
Thrust
Thrust
0
- 5
dFSM &
Stress (MPa)
550
550
300
6
4
2
0
- 2
PrinZ
dFSM
450
Angermanland
Sweden
Vesteralen
Norway
350
500
Hmin
Hmin
450
400
Hmax
Hmax
500
400
PrinZ
300
6
3
0
- 3
- 6
Thrust
dFSM
Thrust
350
550
550
Hmax
Hmin
500
450
400
500
Hmin
450
Helsinki
Finnland
Gallivare
Sweden
400
PrinZ
350
300
1 0
5
0
- 5
Hmax
350
PrinZ
dFSM
- 1 5
dFSM
Thrust
- 1 0
- 5
300
5
Thrust
0
0
Time (ka BP)
Figure 7b. Same as Fig.7a except for six sites in Fig.2b.
- 1 5
- 1 0
- 5
Time (ka BP)
0
onset time around 11-9 ka BP and maximum instability is reached around 10-7 ka BP. Thus, the
predicted onset time of Gallivare coincides with the observed time of formation of the nearby postglacial
faults [6, 26]. Again the mode of failure is predicted to be thrusting and this is precisely what is
observed for the postglacial faults. However, the mode of failure of current earthquakes is not restricted to thrusting [6, 7] - indicating that tectonic stress may have played a comparatively more
significant role today than at postglacial times (due to the relaxation of the rebound stress). Inspecting
the magnitudes of instability, it can be seen that in Angermanland and Oslo, dFSM were as low as -4
MPa and in Gallivare -3 MPa was also achieved during early postglacial time. These predicted magnitude are at least twice as large as those in Laurentia during postglacial time. If these values of fault
instability is an indicator of the magnitude of rebound stress available to trigger earthquakes, then this
larger magnitude of dFSM may result in a larger throw of the postglacial faults in Fennoscandia
provided that rock friction for fault reactivation in Laurentia is comparable if not higher than that in
Fennoscandia. This may partly account for the much smaller throw seen in Laurentia.
Figure 7 also shows that the compressive stress within the ice margin generally decreases in
amplitude from glacial maximum to the present (e.g. Melville Pen., Ungava Pen., Charlevoix, Oslo,
Orebro, Angermanland, Gallivare and Helsinki). However, for sites near the ice margin (Cartwright,
Newfoundland and Vesteralen), the compressive horizontal stress increases slightly from around 9 ka
BP to the present. This is due to the inward migration of the peripheral bulge characteristics of deep
flow models.
Fig.8 shows the change in orientation and magnitude of the horizontal principal stress from 9
ka BP to the present if tectonic stress difference (SHmax-Shmin) is 5 MPa. At 9 ka BP, the average
orientation of the maximum principal stress (thick lines in Fig.8a) in E. Canada is along the ENE
direction - however stress orientations change by about 30 degrees as one goes from the Great Lakes
to northern Quebec and Labrador. Note that for the land grid J8 and in some ocean grids along the
Baffin Bay-Labrador Sea, stress orientation is mainly in the NW-SE direction. This is because rebound stress difference is generally smaller than the tectonic stress difference (5 MPa) except at the
land grid J8 and at the ocean grids along the Baffin Bay-Labrador Sea, therefore rebound stress is able
to dominate the stress orientations in these special places. However, by the time at 0 ka BP (Fig.8c),
rebound stress difference in all the land grids (including grid J8) have decayed below the 5 MPa level
- so that all orientations became dominated by the tectonic stress and uniform stress orientations
(within 10 degrees of ENE) are predicted on land. For the ocean grids along the Baffin Bay-Labrador
Sea, rebound stress difference has also decayed - but not enough for the tectonic stress to be dominant,
thus the orientations offshore remain NW-SE to E-W.
It should be noted that the predicted orientation in grid J8 at 9 ka BP is close to the overall
stress orientations inferred from the postglacial faults in southeastern Canada, except that the location
of J8 is too far to the north. There are two reasons for this discrepancy: First of all, stress orientations
within the ice margin are strongly perturbed by the presence of local ice domes which are not resolved
by the spatial resolution of the ICE3G model. Secondly, the dimensions of the finite element grid,
which is limited by the spatial resolution of the ice model, result in a coarse grid and a reduction of the
magnitude of the rebound stresses near the ice margin (due to spatial averaging in this area of high
stress gradient). Therefore, to explain the stress orientations at 9 ka BP, an ice model with finer spatial
resolution is required. Due to this limitation, we will not compare the predicted and the observed
paleostress orientations on land, but we will continue to examin if the earth model can explain the
observed rotation of stress orientations in southeastern Canada during the last 9,000 years.
0E
20E
0W
12
W
40E
60
10
W
1
1
2
2
3
60
3
4
N
4
5
5
6
6
0E
7
0W
7
12
8
8
9
60N
9
40
10
10
N
(a) TIME
11
Max Stress = 190.8 MPa
A
B
C
D
11
Max Stress Magnitude = 157.7 MPa
Time = 9 KBP
= 9 ka BP
E
F
12
(b)
12
G
H
I
J
K
L
M
N
A
65
B
C
D
E
F
G
H
I
J
K
3
4
3
5
4
6
5
7
0E
2
6
8
7
9
8
10
9
11
60N
12
TIME = 0 ka BP
11
12
(d) Time
= 0 KBP
Max. Stress Magnitude = 145.7 MPa
80W
100W
Max Stress = 186.8 MPa
20E
(c)
10E
N
40E
10
40
30E
12
0W
N
L
2
N
1
60
1
Figure 8. Orientation and magnitude of the projected horizontal principal stresses 9000 years ago
(a,b) and at present (c,d). This is calculated for model L1 where the tectonic stress has a
differential magnitude of 5 MPa.
On the other hand, the predicted stress orientations offshore the Canadian east coast do not
suffer from the lack of spatial resolution of the load because the spatial variation of the ocean load is
much more gradual. Comparing the stress orientations predicted offshore at 9 ka BP with the current
observed orientations (Fig.2a) shows remarkable agreement between them, suggesting that the anomalous stress orientations offshore may be caused by postglacial rebound. However, the predicted
orientations offshore at the present (0 ka BP) do not show such good agreement with the observations
today, indicating that the viscosity in the lower mantle may not be high enough. This disagreement
may also be due to the fact that stress orientations offshore are not well determined [9]. For this
reason, stress orientations offshore will not be considered any further.
In Northern Europe, similar rotations in stress orientations are also predicted during the last
9,000 years. At 9 ka BP (Fig. 8b), the orientations of the maximum principal stress is predicted to be
non-uniform in Fennoscandia: ranging from E-W and WNW in northern Sweden to NW near the
centre of rebound and again to E-W in northern Germany and then back to NW far from the southern
edge of the ice sheet. However, at the present, stress orientations are more uniform and are predicted
to lie within 10 degrees of the WNW direction (Fig. 8d). The areas onland that experience most
significant stress rotation are predicted to be the area northeast of Leningrad and in southwestern
Norway. The amount of rotation is smallest near the center of rebound around Angermanland, thus,
the paleostress orientations inferred from the postglacial faults near Gallivare (Fig.2b) is consistent
with the predicted at 9 ka BP and at the present [11, 42, 43]. The reason is because SHmax due to ridge
push at the North Atlantic is also aligned approximately in the same (NW) direction.
It should be noted that Fig. 8 is computed by assuming a 5 MPa tectonic stress difference. In
general, stress orientations depend on the relative magnitude between the horizontal tectonic stress
difference (SHmax-Shmin) and the rebound stress difference [50]. However, with given ice and earth
models, the relaxation of rebound stress difference is determined. But, tectonic stress difference is
poorly known, thus a range of values will be used below to study its effect on the stress orientations at
different sites in E. Canada and N. Europe. The location of the sites are given by the grid names in
Fig.8.
Figs. 9a & b show that, in North America, with (SHmax-Shmin) in the range 4-10 MPa, the
predicted stress orientations on land are non-uniform at 9 ka BP but rather uniform (within 15 degrees) at the present. For the site J8, stress orientation rotates about 90 degrees during the last 9000
years if the tectonic stress difference is 4 MPa. However, if the tectonic stress difference is larger than
10 MPa, then the temporal stress rotations are smaller (<10 degrees) because the orientation of the
total stress field becomes dominated by the static tectonic stress. On the other hand, for tectonic stress
difference less than about 4 MPa, large stress rotations will occur, but the predicted stress orientations
on land today will be non-uniform - with stress orientation in Labrador (e.g. grid L6) differing from
that in the Great Lakes (e.g. grid H11) by more than 50 degrees if (SHmax-Shmin) is less than 2 MPa.
Thus, in order to explain the contemporary uniform stress orientations and the large stress rotations
since postglacial times, tectonic stress must be in the range 4-10 MPa. These results are in close
agreement with the findings of Adams & Bell [5].
The effect of tectonic stress difference on stress orientations in northern Europe is shown in
Figs.9 c & d. Inspection of Fig.9d shows that when the tectonic stress difference is greater than about
5 MPa, then it determines the contemporary stress orientations and a uniform stress field is obtained at
0 ka BP. However, for tectonic stress difference less than about 4 MPa, then rebound stress will be
able to dominate the contemporary orientations resulting in a non-uniform stress field. At 9 ka BP
(Fig.9c), the transition between uniform and non-uniform stress orientations occurs around 15 MPa.
Thus, if stress orientations had been uniform throughout the last 9,000 years, then the tectonic stress
difference must be at least 15 MPa. However, if the paleostress orientations were non-uniform but
became uniform at the present, then tectonic stress difference must be about 5-15 MPa. Since we have
no information on whether the paleostress field were uniform or non-uniform, the effect of tectonic
stress difference in N. Europe will not be pursued any further.
To conclude this section,let us consider the predicted strain rates over E. Canada and N. Europe since strain rates are also associated with earthquakes [20]. Of the six components of strain, we
shall only consider the vertical components: namely vertical strain rate ε zz and vertical shear strain
rate ε zh = ε zx + ε zy. The former is related to vertical uplift velocity due to the postglacial rebound
process and the latter is related to the shear stress. The spatial distribution of these strain rates at the
present time are contoured in Fig.10. Inspection of Figs.10a & c shows that vertical strain rates peak
near the center of rebound where vertical velocity is the highest. The estimated peak value of -7 x
10 –17 (1/sec) in E. Canada( or -2.2 x 10 –9 /year) is close to what James & Bent [20] found. Inspection
of Figs.10b & d shows that vertical shear strain rate has peaks and troughs closer to the ice margin at
Orientation of Max Principal Stress
(degrees anticlockwise from East)
For sites in E. Canada
120
120
a) t = 9 KBP
b) t = 0 KBP
H11
100
100
J10
80
J8
80
60
L6
60
40
40
20
20
0
0
0
10
20
30 0
10
20
30
Orientation of
Max Principal Stress
(degrees clockwise from East)
For sites in N. Europe
120
120
d) t = 0 KBP
c) t = 9 KBP
C6
100
100
G3
80
80
G10
60
H7
60
40
40
20
0
20
-20
0
0
5
10
15
20 0
5
10
15
20
Difference between Horizontal
Tectonic Principal Stress (MPa)
Figure 9. The effect of tectonic stress difference on the orientation of SHmax is shown for 4 locations
in North America(a, b) and 4 others in Northern Europe (c,d). Results are computed with
Model L1.
the last glacial maxima. It is of interest to note that in E. Canada, large earthquakes are found in the
trough of shear strain-rate along the Baffin Bay-Labrador coast (Figs. 10b & 1a). Clusters of large
earthquakes are also found near the peaks in Southampton Island (north of Hudson Bay) and the
mouth of the St. Lawrence river (Figs. 10b & 1a). In N. Europe, clusters of earthquakes are also found
in the peak of shear strain-rate in southern Norway around Oslo and Berger (Figs. 10d & 1b). The
same can be said about the peak in the northwest coast of Norway around Vesteralen. However, other
peaks in Fig.10b & d are not associated with any clusters of earthquakes and there are clusters of
earthquakes in Fig.1a & b which do not correspond to any peak or trough in Fig.10b & d. Thus, the
A) Vertical Strain Rate at t=0 ka BP
(1/sec) x 1E17
C) Vertical Strain Rate at t=0 ka BP
(1/sec) x 1E17
-1
0
6
0
3
-2
-2
60
0
N
-1
-1
-5
N -3 -4
-6
3
-4
-6
-3
-5
-3
0
1
-9
-9
-6
-2
-6
0
-6
0W
-4
-2
12
-6
-7
-9
-6
-7
-4
-3
-3
3
4
-1
0N
60
11
N
0
W
0
B) Vertical Shear Strain Rate at t=0 ka BP
( 1/sec) x 1E17
0
0
2
-2
-4
0
N
2
-1
6
-2
D) Vertical Shear Strain Rate at t=0 ka BP
(1/sec) x 1E17
0
-2
-1
3
-2
0
-2
0
0
-5
0
-3
-6
-7
-3
-1
40
-4
-5
-4
60
1
N
1
-2
-2
-1
-1
0W
-3
-2
0
4
0
12
-2
0
-1
-3
6
1
-2
-2
0
0
0
4
6
2
2
4
1
0
0
2
0
2
0
0
N
2
40
N
60
40
-2
0
1
W
0
Figure 10. Spatial distribution of current strain rates predicted over E. Canada and N. Europe. Vertical
strain rates (ε zz ) with units of [1/sec] is multiplied by 10 17 before being plotted in (a) and
(c). Vertical shear strain rates (ε zh ) with units of [1/sec] is multiplied by 10 17 before being
plotted in (b) and (d).
correlation between the location of some peaks or troughs in vertical shear strain rate maps and the
spatial distrubution of current earthquakes is likely to be purely fortuitous indicating that intraplate
earthquakes involves more than just vertical shear strain rates. Since we have shown earlier in this
section that the spatial distribution of current seismicity can be explained by taking the initial Fault
Stability Margin to be greater than 1 MPa everywhere except at the pre-weakened zones in E. Canada
and coastal Fennoscandia, where it is set at zero, strain rates will not be discussed any further.
In summary, the observed spatial distribution and mode of failure of earthquakes, their onset
time, the observed current stress orientations in Eastern Canada and Northern Europe and the rotation
of stress observed in E. Canada can all be explained by this earth model. The only exceptions are some
modes of failure in the northeast coast of Baffin Island and several sites in eastern United States and
Fennoscandia, where local tectonic stress may have played a slightly more dominant role than the
contemporary rebound stress.
IV.b. Effect of High viscosity Lower Mantle
In the last subsection, we saw that changes in the magnitude of the rebound stress trigger
postglacial faulting/earthquake and cause rotation of stress orientations. However, the rate of relaxation of rebound stress is determined by the viscosity structure of the mantle, thus, it is important to
investigate how the results of the last section will be affected if the earth has a different viscosity
profile. This is particularly important because recent investigations of postglacial sea-levels and Earth
rotation point to an increase in viscosity across 670 km and 1200 km depth in the lower mantle [14,
31, 32, 33, 34]. Furthermore, Spada et al. [40] claimed that only a high viscosity lower mantle is able
to trigger earthquakes in E. Canada and N. Europe. However, in their treatment of fault potential, only
rebound stress difference was considered, and the effects of mean stress and tectonic stress have been
neglected (see equation 1). In this section, dFSM as defined in equation (1) will be used to study the
dFSM at 18 ka BP
60
2
6
0
2
64
10
N
10
16
12
0W
9
6
6
7
7
12
2
6
0
6
4
4
2
0N
1
W
0
dFSM at 0 ka BP
dFSM at 0 ka BP
8
2
6
4
-2
-2
1
-2.0
-3.5
-1.0
1
-1.5
-2.5
-4.0
-4
1
-3
-2
0
0W
-2.5
12
-5
-1.0
-5
-3
-1 -2
-3.5
-1.0
-3.0
-3.0
-0.5
-4
-4.0
-2.0
0.0
-1
0
-1.0
-1.5
2
N
1
0.0
-0.5
-3.0
0
-3
0.5
N
2
0
1
6
0
-1
1
60
2
6
6
2
3
8
10
4
4
12
14
0
5
7
8
-2.0
-4
-2
-3
40
-1
0
N
0.0
-0.5
60
N
3
6
8
16
2
40
6
8
16
14
12
60
7
8
14
8
N
3
10
6
40
1
1
4
18
8
5
8
6
7
0
6
4
5
2
0
2
12
3
4
N
8
4
8
6
1
10
6
8
4
dFSM at 18 ka BP
W
Figure 11. Similar to Fig.6 except for a high viscosity lower mantle (model L2) at two time steps.
effects of a high viscosity (1022 Pa-s) lower mantle on the onset time of instability, mode of failure,
magnitude of instability, stress orientations and rotation.
Fig.11 shows that a 1022 Pa-s lower mantle gives larger range of values for dFSM - with higher
peaks but lower troughs both in E. Canada and N. Europe when compared with Fig.6. In particular, the
magnitude of the troughs at 0 ka BP go from -1 MPa to about -4 MPa in both Laurentia and Fennoscandia.
Fig.12 shows more clearly the temporal variation of dFSM, the stresses and the mode of failure.
Comparing Fig.12 with Fig.7 shows that the mode of failure remains the same, while the increase in
the magnitude of dFSM generally results in an insignificant shift in the onset time of instability. The
largest change can be found for sites such as Newfoundland where the small changes in the values of
600
550
500
Melville
600
Hmax
Hmin
Hmax
Hmin
Pen.
550
Ungava Pen.
450
500
450
PrinZ
PrinZ
400
400
1 0
5
0
- 5
dFSM
dFSM
Thrust
1 0
5
Thrust
0
- 5
600
Hmax
550
500
Hmax
Hmin
Hmin
Labrador Sea
near Cartwright
550
500
NW NewFoundland
450
450
PrinZ
400
2
1
0
- 1
dFSM
dFSM
Thrust
1
600
Hmax
Hmax
Hmin
550
550
Hmin
500
Charlevoix, Que.
Indiana
450
450
PrinZ
PrinZ
400
1 0
5
0
- 5
400
2
0
600
500
PrinZ
Thrust
dFSM &
Stress (MPa)
600
400
dFSM
- 1 5
dFSM
Thrust
- 1 0
Time
- 5
(ka BP)
5
Thrust
0
0
- 1 5
- 1 0
- 5
Time (ka BP)
Figure 12a. Similar to Fig.7a except for a high viscosity lower mantle (model L2).
0
dFSM result in a much delayed onset of fault instability. This suggest that accurate determination of
the timing of earthquakes near the ice margin may be more useful for constraining mantle viscosities
than for sites near the center of rebound.
Fig.12b also shows that, starting from the last deglaciation, maximum instability in Fennoscandia
is attained around 8 ka BP with a value of about -6 MPa. However, for Laurentia, fault instability has
been increasing since deglaciation, reaching the value of -3 MPa at present - although the rate of
increase is much slower now. The difference in maximum instability attained between Laurentia and
Fennoscandia again implies that larger rebound stress is available to reactivate the faults and, as discussed in the last section, one can argue for larger throw for the postglacial faults in Fennoscandia
550
550
500
450
400
Hmax
Hmax
Hmin
Hmin
Oslo,
Norway
450
Orebro,
S. Sweden
400
PrinZ
PrinZ
350
300
8
4
0
- 4
500
dFSM
dFSM
Thrust
350
300
1 0
5
Thrust
0
- 5
dFSM &
Stress (MPa)
550
550
Hmax
500
Hmin
450
400
500
450
Angermanland
Sweden
Vesteralen
Norway
PrinZ
350
300
6
4
2
0
- 2
Hmax
400
PrinZ
350
300
5
Thrust
dFSM
Thrust
0
dFSM
- 5
550
550
500
450
400
Hmax
500
Hmin
Hmin
450
Helsinki
Finnland
Gallivare
Sweden
400
PrinZ
350
300
1 0
5
0
- 5
Hmax
PrinZ
350
300
dFSM
- 1 5
dFSM
Thrust
- 1 0
- 5
5
Thrust
0
0
Time from the present (KBP)
- 1 5
- 1 0
- 5
0
Time from the present (KBP)
Figure 12b. Similar to Fig.7b except for a high viscosity lower mantle (model L2).
-0.06
-0.09
0.00
0.06
-0.03
0
6
-0.06
-0.00
60
-0.06
N
-0.03
-0.03
0.18
-0.06
-0.12
N
0W
-0.00
0.12
-0.09
-0.09
-0.06
-0.00
0.06
-0.06
0.24
0.18
12
0.03
0.03
-0.00
0.00
0.18
0.09
0.00
0.06
-0.03
0.12
0.24
0.06
0.06
0.06
-0.06
-0.09
-0.12
0.00
-0.00
40
-0.09
-0.09
N
-0.06
W
-0.03
N
60
40
0.12
0.09
20E
0E
12
W
10
W
1
40E
60
0W
Figure 13. Spatial distribution of the present rate of change for dFSM in Laurentia and Fennoscandia
for Model L2 at a depth of 12.5 km. Contours are in MPa/ka. Solid contour lines indicate
decreasing fault instability while dashed contours indicate increasing fault instability.
65
2
N
3
60
N
4
65N
5
6
0E
0W
7
12
8
9
60N
10
N
TIME = 9 ka BP
Max Stress = 186.5 MPa
A B C D
60N
11
E
F
TIME = 9 ka BP
12
G
H
I
J
K
L M N
A
D
E
F
G
H
I
J
K
L
10
1
B C
Max Stress=180 MPa
W
40
2
60
65
N
3
N
4
65N
5
0W
7
0E
6
12
8
9
60N
10
TIME = 0 ka BP
12
TIME = 0 ka BP
Max Stress=169 MPa
40E
80W
Figure 14. Similar to Fig.8 except for a high viscosity lower mantle (model L2).
30E
Max Stress = 180 MPa
100W
60N
11
20E
N
10E
40
Orientation of Max Principal Stress
(degrees anticlockwise from East)
than in Laurentia. That fault instability in Laurentia is predicted to increase from the end of deglaciation
implies that earthquake magnitude and frequency in E. Canada will increase in the future, moreover,
previously stable area may become more earthquake prone. This is particularly important for the
planning of nuclear waste repositories which need to be located in areas that stay stable for the next
few thousand years. Fig. 13 shows the spatial distribution of the present rate of change for dFSM.
Areas with solid contours are sites where the magnitude of rebound stress available to trigger earthquakes is decreasing. However, areas with dashed contours are areas where fault instability is increasing and thus not suitable sites for the safe storage of nuclear waste in the next few thousand years if the
viscosity of the lower mantle is 1022 Pa-s. However, if the lower mantle viscosity is 1021 Pa-s (Model
1), then fault instability will be decreasing in the deglaciated areas and this would not be a problem.
For sites in E. Canada
120
120
a) t = 9 KBP
b) t = 0 KBP
H11
100
100
J10
80
J8
80
60
L6
60
40
40
20
20
0
0
0
10
30 0
20
10
20
30
Orientation of
Max Principal Stress
(degrees clockwise from East)
For sites in N. Europe
120
120
t = 9 KBP
C6
t = 0 KBP
100
G3
100
G10
80
80
H7
60
60
40
40
20
20
0
0
0
5
10
15
20 0
5
10
Difference between Horizontal
Tectonic Principal Stress (MPa)
Figure 15. Similar to Fig. 9 except for a high viscosity lower mantle (model L2).
15
20
The increase in lower mantle viscosity also has important effects on the contemporary stress
orientation. Due to the longer relaxation times, rebound stress difference remains large even at the
present and for a 5 MPa tectonic stress difference, rebound stress will dominate the stress orientations
even up to now, resulting in little stress rotation. Thus the predicted stress orientations remain nonuniform in Laurentia and Fennoscandia throughout the last 9,000 years (Fig.14). Since non-uniform
stress orientations are not consistent with the observations today, the tectonic stress difference must be
at least 18 MPa in Laurentia and Fennoscandia (Fig.15) so that a uniform stress orientation field is
obtained for earth model L2. However, with such high tectonic stress difference, the observed rotation
of stress orientations in Laurentia cannot be explained.
In summary, a high viscosity lower mantle will have no effect on the mode of failure. The
effect on the onset time of instability is insignificant except for sites near the ocean margin, however,
even there, more accurate age determination is needed before it can be used to infer lower mantle
viscosity. On the other hand, a high viscosity lower mantle can result in areas in E. Canada where fault
instability will increase in the next few thousand years and thus not suitable for the safe storage of
nuclear waste. Finally, stress rotation is very sensitive to the viscosity of the lower mantle. With
viscosity as high as 1022 Pa-s in the lower mantle, any significant rotation of stress orientation during
the last 9,000 year is prevented, thus it is not possible to explain the observed stress rotation and
contemporary stress orientations in E. Canada simultaneously.
IV. Conclusions
In conclusion, we have seen that although glacial unloading gives fault instability of the order
of a few MPa which is probably not large enough to dictate fracture, nevertheless intraplate earthquakes in E. Canada and N. Europe can still be triggered by postglacial rebound stress through reactivation of optimally oriented pre-existing faults in tectonically pre-weakened zones. This resolves the
issue as to whether tectonic stress or postglacial rebound stress is more important in the generation of
intraplate seismicity in Laurentia and Fennoscandia. We have also shown that the spatial distribution
of earthquakes, their onset time, the observed current stress orientations in Eastern Canada and Northern
Europe, the rotation of stress observed in E. Canada and most of the observed modes of failure can all
be explained simultaneously if the mantle has a uniform 1x1021 Pa-s mantle.
The effect of a 1022 Pa-s lower mantle on the ability to explain the observations has also been
investigated. It is found that the magnitude of stress, dFSM and the timing of the onset of fault instability are sensitive to viscosity structure. Unfortunately, direct stress measurements are few and dFSM
cannot be observed. Also, changes in dFSM is also small, thus, the role of rebound stress is still
limited to the reactivation of pre-existing faults generated by previous tectonic events - i.e. both tectonic forces and rebound stress are still needed to explain current earthquakes in E. Canada (evidence
in the distribution of contemporary earthquakes). As for the timing of the onset of earthquakes, their
accuracy has not been determined well enough for the inference of mantle viscosity. Thus, until higher
accuracy in the determination of the onset of earthquakes/faulting events is available, the proposal of
Spada et al. [40] to constrain mantle rheology from seismicity cannot be realized.
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