Earthquakes and geological structures of the St. Lawrence Rift System - Intraplate Earthquakes - page 81

Since the arrival of the first Europeans in the early 1600s, the CSZ has been subject to five
earthquakes of magnitude 6 or larger: in 1663 (M 7), 1791 (M 6), 1860 (M 6), 1870 (M 6

½ ),

±

and 1925 (magnitude M S 6.2 0.3) (Lamontagne et al ., 2008a ) . The earthquake potential

of the area led the Government of Canada to conduct two field surveys that defined its
main seismotectonic characteristics (Leblanc et al ., 1973 ; Leblanc and Buchbinder, 1977 ) .
Earthquakes occur between the surface and 30 km depth, in the Precambrian Shield,
which outcrops on the north shore of the St. Lawrence River or is found beneath Logan's
Line and the Appalachian rocks ( Figure 4.3 ) . Hypocentres cluster along or between the
mapped Iapetan faults (also called St. Lawrence paleo-rift faults). The largest earthquake of
the twentieth century was the 1925 earthquake, and its focal mechanism has one nodal
plane consistent with a reactivation of a southeast-dipping paleo-rift fault (Bent, 1992 ) . The
installation of a permanent seismograph network in 1978 ( Figure 4.3 ) has helped to define
additional characteristics of the area. The St.-Laurent fault, one of the major rift faults of
the CSZ, was formed in the late Precambrian but was also active after the Devonian meteor
impact (Rondot, 1979 ) , probably during the early stages of the opening of the Atlantic
Ocean in late Triassic-Jurassic times (Lemieux et al ., 2003 ) . This fault is not particularly
active but appears to bound concentrations of hypocentres (Lamontagne, 1999 ) .
Due to its concentration of earthquakes, the CSZ has been the focus of various geophys-
ical studies (Buchbinder et al ., 1988 ) . Investigations of velocity structure include a seismic
reflection-refraction survey (Lyons et al ., 1980 ) , microearthquake surveys (Leblanc et al .,
1973 ; Leblanc and Buchbinder, 1977 ; Lamontagne and Ranalli, 1997 ) , analysis of teleseis-
mic events (Hearty et al ., 1977 ) , receiver function analysis (Cassidy, 1995 ) , shear wave split-
ting and anisotropy studies (Buchbinder, 1989 ) , and focal mechanisms of microearthquakes
(Lamontagne, 1998 ; Bent et al ., 2003 ) . The concentration of earthquakes led to earthquake
prediction studies in the late 1970s and early 1980s (Buchbinder et al ., 1988 ) . Roughly
80% of Charlevoix earthquakes occur in the depth range 5-15 km in Grenvillian basement
rocks, with some as deep as 30 km ( Figure 4.3b ) . Comparing this depth distribution to
rheological models of the region, Lamontagne and Ranalli ( 1996 ) attribute earthquakes to
faulting above the brittle-ductile transition to depths of at least 25 km. The reactivation
of pre-existing faults could be due to high pore-fluid pressure at temperatures below the
onset of ductility for hydrated feldspar at about 350 C and/or a low coefficient of friction,

possibly related to unhealed zones of intense fracturing. The distribution of spatially clus-
tered earthquakes within the Charlevoix seismic zone indicates that very few earthquakes
have occurred on the same fractures with similar focal mechanisms, implying that these
fault zones occur in highly fractured rocks, especially those within the boundaries of the
Devonian impact structure (Lamontagne and Ranalli, 1997 ; Figure 4.3 ) . The hypocentre-
velocity simultaneous inversion of local P and S waves produced a velocity model that
revealed areas of high-velocity bodies at mid-crustal depths (Vlahovic et al ., 2003 ; Figure
4.3b ) . These areas were interpreted to be stronger, more competent crust that separates
CSZ earthquakes into two main bands elongated along the St. Lawrence River. Mazzotti
and Townend ( 2010 ) noted that this seismic zone contains evidence for a local rotation of
S Hmax (direction of maximum horizontal stress axis) from the regionally NE-SW-oriented

°