Crustal faults located close to cities may induce catastrophic damages. When
recurrence times are in the range of 1000–10 000 or higher, actions to
mitigate the effects of the associated earthquake are hampered by the lack of
a full seismic record, and in many cases, also of geological evidences. In
order to characterize the fault behavior and its effects, we propose three
different already-developed time-integration methodologies to define the most
likely scenarios of rupture, and then to quantify the hazard with an
empirical equation of peak ground acceleration (PGA). We consider the
following methodologies: (1) stream gradient and (2) sinuosity indexes to
estimate fault-related topographic effects, and (3) gravity profiles across
the fault to identify the fault scarp in the basement. We chose the San
Ramón Fault on which to apply these methodologies. It is a
In active margins, sustainable balance between city development and
geological environment requires understanding seismic hazard to reduce
the associated risks. When city emplacements are adjacent to potentially
active crustal faults, seismic risk are elevated; thus, it is important to
quantify their possible effects. A good example to demonstrate the potential
danger of crustal faults is the Chūetsu earthquake of
Geological map of the zone (Thiele, 1980; Fernández, 2003). In the figure the location of the Apoquindo hill outcrop (Fig. 3c) and the TEM profile can be observed (Fig. 3a).
Defining whether or not the fault is active is crucial in affirming that the
risk exists. An active fault that is preferentially oriented with respect to
the current tectonic regime allows stress release, eventually triggering
earthquakes (examples of preferentially oriented faults are normal and trust
faults subject to Andersonian stress regime (Anderson, 1951), with a strike
perpendicular to
Along the SRF, the probable low recurrence of characteristic earthquakes,
the stream gradient index, which can compare the relative uplift rate in
a certain area by studying the topographic profile of near-fault rivers (e.g.,
Font et al., 2010; Casa et al., 2010); the sinuosity index, which estimates the
uplift of a fault by observing the sinuosity of the mountain front (Bull and
McFadden, 1977); across-fault gravity profiles to estimate the shape of the
fault scarp in the basement beneath the sedimentary basin, given the large
density contrast between rocks and sediments, and also because basement
morphology is a useful marker of cumulative faulting. Since the SRF has a low slip rate, fault scarp morphology may be modified by deposit and/or erosion
surface processes. Thus, we favor the use of gravity profiles and
geomorphological measurements instead of scarp topographic analyses. To
develop the geomorphological methodologies, we used a 30 m resolution DEM
(SRTM30; Farr et al., 2007), and for gravity we used a Scintrex CG-5 Autograv gravity meter and an R4 Trimble DGPS.
As a final product of this research we tackle the difficult question of
relating the occurrence of a given seismic event with an associated damage
prediction. One possibility is to estimate the expected acceleration during a
characteristic earthquake, and then link this output with damage.
Acceleration is an objective measure of the seismic effects, and thus it is
not affected by the quality of the housing or infrastructure. We chose
empirical equations for crustal earthquakes (e.g., Sadigh et al., 1997; Chiou
and Youngs, 2014) to predict the peak ground acceleration (PGA). The
robustness of this methodology is grounded in the last decade of understanding
of the key variables that control the PGA. Principal variables are event
magnitude, fault type, hanging wall, and site effects (near field effects). We
chose the Chiou and Young equation (2014) because this model also accounts
for a low slip rate crustal fault, and has an extensive record of different
earthquakes worldwide. In order to use this particular approach, further
parameters are required, such as the shallow depth of the rupture and the dip
of the fault. We estimated these parameters with a joint interpretation of
the surface geology (Armijo et al., 2010; Rauld, 2011; Vargas et al., 2014),
the results of the microseismic study, and a high-resolution 2-D
geo-electrical TEM study.
This region has been dominated by the subduction of the Nazca Plate underneath the
South American Plate since at least the Jurassic time (Mpodozis and Ramos,
1989). Upper basement rocks are dominated by the volcano–sedimentary Abanico
and Farellones formations. These volcano–sedimentary sequences are mainly
constituted by pyroclastic and lava strata, interdigitated with lava and
different sedimentary rocks. The earliest Abanico Formation was deposited
during the late Eocene–Oligocene and was reformed later on (Charrier et
al., 2002; Godoy et al., 1999). The Farellones Formation was
deposited above the underlying sequence during the early and middle Miocene
(Charrier et al., 2002). These volcano–sedimentary sequences are also
intruded upon by Miocene plutons (Kurtz et al., 1997; Thiele, 1980). Among them,
we find the La Obra pluton 19.6
The Santiago Basin sediments can be grouped into four main sequences. The
widespread and well-compacted fluvial sediments, associated with the
material transport along the Maipo and Mapocho rivers (i.e., Leyton, 2010;
Yañez et al., 2015). In addition there are the alluvial and colluvial
deposits, which are semi-compacted and spatially concentrated in the piedmont
of the Andes (Fernández, 2003). Finally, and restricted to the northern
area of the study, it is possible to find fine soil, mostly lacustrine
deposits, and to a lesser extent in the western and southern area of the
study, pyroclastic ash related to the Maipo volcanic eruption
The San Ramón Fault has been studied using high-resolution DEM,
satellite images, and field observations (Armijo et al., 2010; Rauld, 2011).
These authors conclude that the SRF is a west-verging reverse fault that
accommodates the compressive stress regime in this segment of the central
Andes (see Fig. 1). One of the best outcrops from which to observe the fault can be
examined in Fig. 3c. The ignimbrite of the Maipo eruption
In order to get an estimate of the seismic hazard of the SRF, we consider five steps (see flow chart in Fig. 2). In Sect. 3 we briefly describe these methodological steps, and in Sect. 4 we present the results derived from their application.
To achieve the first goal, we deployed a small seismic network of five
borehole seismometers with three-component 2 Hz sensors (short period
S31f-2.0a of IESE) running in continuous mode during a 1-year time window
with a sample rate of 100 Hz. The equipment was installed near the fault
trace, covering an area of 20 km by 15 km (see Fig. 4). The preliminary
estimate of the origin times and hypocentral coordinates was determined by
means of the HYPOINVERSE program (Klein, 1984), considering the initial
velocity model proposed by Villegas (2012) for this area. Then we made a
recursive process for the best estimate of velocity structure and event
location using the VELEST and HYPOINVERSE code (Kissling et al., 1995). The
model with the lowest RMS error has 10 layers and assumes a varying
Scheme of objectives and methodologies used in this work. The final objective is in yellow.
The association between seismic events and the SRF is determined by the following procedure. Given the SRF surface trace we project its potential extension downwards using an empirical relation (Wells and Coppersmith, 1994). Well-localized events that are located inside the area of fault influence are considered a likely representation of fault activity. These events were projected onto a central cross section perpendicular to the fault for visual discrimination of truly fault-related events (see red and blue rectangle in Fig. 4). Comparing the number of events related to the fault with those from other structures, the importance of the fault in the stress release of the whole zone can be discussed. Examples of other structures in the study area are the El Diablo fault, Chacayes back-thrust (Farias et al., 2008), the unnamed faults that fold the Abanico and Farellones units (Armijo et al., 2010; Rauld, 2011), and the structure associated with the “Santa Rosa cluster” (Leyton et al., 2009).
To determine the geometrical characteristics of the fault, we consider the data available in the literature (Armijo et al., 2010; Vargas et al., 2014),
the spatial distribution of seismic events, and a 2-D resistivity image of a
well-maintained scarp. The electrical imaging was obtained by carrying out a
high-resolution TEM (transient electro magnetic) experiment that provides a
good constraint for the first 150–200 m fault section. TEM technique is a
geophysical method for obtaining an electrical resistivity image of the
subsurface (for details of TEM theory and data processing see
Telford et al., 1990, for instance). In the field, we used the FastSnap TEM System, completing
24 TEM stations 25 m apart. To observe the location
of the TEM profile see Fig. S1 in the Supplement. For the
purpose of getting maximum spatial resolution we use in-loop
configuration with a transmitter (Tx) loop of 25
An empirical first-order relationship between rupture length and earthquake
size (Blaser et al., 2010; Wells and Coppersmith, 1994) allows the estimation
of the characteristic earthquake magnitude based on a well-constrained
rupture length. In this case we use three different methodologies to quantify
this length in terms of the associated uplift. Each independent methodology
estimates the fault uplift, measuring physical or geomorphologic properties.
These methodologies are
across-strike gravimetric profiles. Due to the low erosion rate of the
basement, tectonic deformation is better preserved in the basement compared to
the surface. Density contrast between gravel material and basement rocks
makes the gravity method a suitable tool for estimating basement geometry across
the fault ( stream gradient index. This index represents a relative measure of the
surface uplift based on drain topographic profiles (Hack, 1973; Merritts and
Vincent, 1989). Zones with high values suggest larger surface uplift relative
to zones with low values (e.g., Casa et al., 2010; Font et al., 2010). This
methodology is useful because drainage profiles are good indicators of the
long-term uplift process. To determine the corresponding drainage, we used
the ArcGIS utilities flow direction, flow accumulation, and watershed. We only
chose secondary drainage with similar length to avoid potential bias
associated with different scale processes. To calculate the stream
gradient (hereafter SL) index we separate the topographic profiles of each drainage
into several segments with 50 m of elevation overlap. For each segment,
SL was calculated by multiplying the slope by the middle distance to the drain
top (Hack, 1973; Merritts and Vincent, 1989). sinuosity index. Long-term activity of a piedmont fault can be inferred
from mountain front sinuosity index (Bull and McFadden, 1977). Low values of
this index indicate a fault-controlled landscape (Bull and McFadden, 1977),
and the minimum value is 1.00. This index was developed for normal faults,
but it has been satisfactorily proven in reverse faults (Casa et al., 2010;
Jain and Verma, 2006; Singh and Tandon, 2007; Wells et al., 1988). At the
transition between mountain front (high slope) and basin (low slope),
across-strike slope differences generate an anomaly angle. We used ArcGIS to
generate a slope map and define the best slope angle that represents the
basement–basin contact, comparing the slope map with the most detailed
geology information (in this case Rauld, 2011). In the present study, this
angle is between 15 and 16
Results from each methodology will be
discussed separately, and then a joint interpretation will be made to define
the rupture length of the characteristic earthquake.
Based on the length of characteristic earthquake rupture and fault geometry,
we estimate the seismic hazard by calculating the corresponding peak ground
acceleration (PGA). In this work we used the attenuation model of Chiou and
Youngs (2014), which is appropriate for crustal earthquakes. The PGA field is calculated
over a grid of 1 km spacing. Each grid point is characterized by the
associated basement depth and weighted average shear wave velocity for the
first 30 m (
The hazardous domains derived from the PGA maps are defined in terms of the
acceleration distribution. In order to gain some insight into the effects of an
SRF event, we compared the expected acceleration estimated in this study with
the observed acceleration during the 2010 Chile earthquake in the Maule Region
TEM results and the Apoquindo hill outcrop.
Microseismic study results. In the left panel is the map with
epicenters of the recorded events. The color represents depth of the hypocenter.
In white are the events with localization errors larger than 8 km. White inverted
triangles are the seismic stations. The red rectangle represents the
70
Over the year of microseismic recording we identified 1666 events within a
radius of 150 km around the network. The majority were located on the
Nazca–South American Plate contact, and only 245 of them were crustal
intraplate earthquakes with a depth above 35 km. Of these
crustal earthquakes, 56 % are related to blasts in mining operations, and only the
remaining 44 % (110 events) are associated with natural sources. Some
events were well registered in only two stations, implying large position
errors. To be certain of earthquake locations, we restricted the position
errors to 8 km in both horizontal and vertical coordinates. Well-recorded crustal events are constrained to an 80
The resistivity imaging (Fig. 3) reveals different domains below the
well-preserved scarp (the location of the TEM profile can be found in
Supplement Fig. S1). Electrical domains are subhorizontal in the first
100–150 m depth and homogeneous at both edges. But in the fault core,
electrical domains are clearly subvertical below 100–150 m depth. We
associate the relative conducting and subhorizontal upper domain with the
sedimentary infill of the basin in the foot hills. Below the sedimentary
cover we associated the electrical high- or low-resistivity domains with
pristine or fracture basement rocks, respectively. In consistency with this
observation, it is possible to separate the electrical image into six different
units (see Fig. 3b):
quaternary high-resistivity sediments at the
colluvial wedge of the scarp (mean resistivity of 144 ohm m). quaternary
relatively dry sediments (45 ohm m). upstream quaternary wet sediments
in the fault hanging wall (25 ohm m). pristine basement rock (> 1000 ohm m). low-resistivity domain associated with fluid
percolation along fractured rocks (1.6 ohm m). In some cases, the resistivity
is less than 1 ohm m, probably due to the presence of highly saline
hydrothermal fluids that use fault planes as conduits. unconsolidated
sediments interbedded with hydrothermal fluids downstream with respect to the
fault (0.5 ohm m).
In terms of the fault geometry, this geoelectrical imaging represents a
family of nearly vertical low and/or high resistivity bodies, interpreted as a
system of high-angle faults reaching the surface. This suggests that the
SRF is a system, not a single fault.
In the upper panel the results of gravity inversion of profile 7 are shown.
In the elevation profile the observed surface scarp is drawn (Rauld, 2011).
The regional tendency is calculated by a first order approximation. The
inverse profile has a vertical exaggeration of
Gravity profile results.
Stream gradient index and relative erosion representation.
Gravity modeling demonstrates that cumulative deformation is better expressed in the basement. This behavior is confirmed in the inversion of gravity profiles that cross an evident fault scarp (e.g., Profile 7 in Fig. 5a). The surface scarp observed in the elevation profile (top of Fig. 5a) is less abrupt than the basement scarp observed in the gravity inversion. Basement scarps are characterized by short wavelength (< 200 m) and relatively large (> 10 m) gravity bodies. In addition, the gravity inversion of Profile 7 identifies three scarps in the basement, whereas only two appear at the surface. The erosion process smooths the surface and dilutes the presence of multiple scarps, either by the retreat of previous scarp or as an effect of diffusive erosion (e.g., Carretier et al., 2002).
To understand the along-strike continuity of the basement scarp, and thus the possible rupture length, we mapped all of the basement scarps identified. One example of this scarp continuity analysis is shown in Fig. 5b. In this case we can observe several scarps in N–S strike continuity. Another important measure obtained in the gravity inversion is the accumulated displacement of each profile, which is a complementary approach to estimate the rupture continuity (see Fig. 6c). The accumulated displacement is estimated by the sum of all the basement scarp heights at each profile. Both approaches consistently show scarp discontinuities. Based on this observation we define four different segments with a mean length of 10 km.
The SL only applies to long wavelengths of the topography, with a cutoff that
depends on the available stream network, which is much wider
than the gravity approach in this particular case. The SL results allow the definition of four main
domains (see Fig. 7a) from north to south. The northern one has a
concentration of high SL values, with a N–S high anomaly (
Lithological differences under the drainage can generate misinterpretation of
the SL index as a direct bedrock uplift indicator. Considering that intrusive
rocks are less erodible than volcanic rocks (
Sinuosity index results. In white the SI values associated with a specific part of the fault named as a section (Sect) in this figure are shown. It is important to note that these sections do not necessarily represent the segments of the fault, they are just a discrete separation along the SRF related to changes in sinuosity in the mountain front. As a reference, the geological map of the zone (Fig. 1) is included in the back.
Using the sinuosity of the mountain front, seven sections were defined, whose SI values are summarized in Fig. 8. The results of section 6 will not be considered in this analysis because the fault is located outside the piedmont in this section, and thus the methodology is not valid there. Sections 1, 2, and 4 show values close to 1 (1.17–1.43), as observed in active reverse faults (1.00–1.50) (Casa et al., 2010; Jain and Verma, 2006; Singh and Tandon, 2007; Wells et al., 1988). The specific sections where SI values are close to 1 also coincide with zones in which gravity profiles suggest fault activity. These sections have a mean value of 1.30, which reflects the limited capacity of the SRF to shape the mountain front, compared to the 1.04 mean value of the piedmont fault in the Himalayas (Jain and Verma, 2006). Sections 3, 5, 6, and 7 have higher SI values (> 2.00). Fluctuating SI values indicate the occurrence of different geomorphologic processes dominating the fault scarp at the surface (i.e., Burbank and Anderson, 2001; Jain and Verma, 2006; Casa et al., 2010).
Interpretation of SRF subdomains or segments. It summarizes the
results of the three methodologies used to define the length of the
characteristic earthquake and the segments interpretation. The
northern part of segment 1 has a high value of stream gradient index; additionally it has
a low SI value and one gravity profile with
The seismic moment magnitude
The results shown in Fig. 10a represent the maximum expected PGA in the study area. At each point we choose the largest PGA value from the corresponding rupture of every SRF subfault. Against intuition, the largest acceleration in the northern segment is observed in the footwall block instead of the hanging wall. This can be explained by the infilling of fine sediments (low shear velocity) with larger site effects compared to the basement rocks at the hanging wall. In the other segments to the south, as expected, the largest PGA is observed in the hanging wall. In segments 2 and 4, parts of the hanging wall are filled with sediments, generating hanging wall and site effects with PGA values above 0.5 g. The largest acceleration of 0.8 g shows up in the area already described.
Seismic events spatially associated with the SRF, suggest that the fault is active. If this inference is correct, their depth distribution shows more affinity with a high-angle fault (Fig. 4). This is consistent with the surface expression of the SRF as described by the Apoquindo outcrop and the TEM profile (Fig. 3c). Although reverse fault optimal orientation is low angle, several examples of normal high-angle faults reactivated as inverse faults have been described in the Andean orogenesis (e.g., Charrier et al., 2002). Another aspect is the importance of the SRF on the whole stress release of the zone in terms of the seismic productivity. The natural seismicity distribution in the study area indicates that just five events can be related to the SRF, representing 12 % of the 41 well-localized events and 5 % of all 110 crustal events regardless of their location error (but still within the area of interest). This denotes that the San Ramón Fault is not the only structure in the deforming cordillera. Nevertheless, it involves a significant hazard given its likely active condition and its proximity to the city.
To estimate the acceleration is necessary to determine some fault first-order
geometrical characteristics. Fault type is the first one, and can generate a
0.8 factor to the acceleration for normal faults and 1.3 for reverse ones,
with respect to strike slip faults (Ambraseys et al., 2005). The TEM profile
and the tilting strata in the Apoquindo hill outcrop clearly demonstrate
reverse kinematics, which is also well supported by many field observations (Armijo et
al., 2010; Rauld, 2011; Vargas et al., 2014). Another relevant variable is
the shallow depth of the rupture (Youngs et al., 1997; Chiou and Youngs, 2008).
In regard to the SRF, the quaternary sediments in the Apoquindo hill are cut by the
fault. Consistent with this observation, TEM results indicate that basement
displacement also reaches the basement roof (Fig. 3), and thus breaks
the surface. Therefore, the shallow depth of rupture is estimated to be at zero
level. The last first-order variable is the fault dip. The NGA-West 2 data
indicate a systematical acceleration increasing with larger dips (Chiou and
Youngs, 2014). The TEM profile and the Apoquindo outcrop show a near-surface
subvertical plane. At depth the microseismic study is also consistent with
a high-angle fault; therefore, the dip angle of the SRF is estimated at 65
According to the integrated analysis carried out the SRF is not necessarily a
continuous fault along its
The northern segment (corresponding to section 1 in Fig. 9) is not necessarily restricted to the length defined in this work because with the information available we cannot trace precise limits. This is mainly because the gravity profiles are a bit sparse. Nevertheless, the existence of the fault in this zone is demonstrated by the stream gradient signal and the gravity profile L2. Despite this drawback, we postulate this segment as a preliminary solution.
In the central area, we propose three segments capable of generating a great earthquake (segments 2, 3 and 4 in Fig. 9) with a high uplift interpreted by the high values of SL and the surface manifestation of the SRF (Armijo et al., 2010; Rauld, 2011). The segmentation of segments 2 and 3 is supported by the gravity profile and the fluctuant sinuosity index and is complemented by lithological changes in the hanging wall unit (see sections 2 and 3 in Fig. 8). In fact, the transition zone between segments 2 and 3 is supported by the lack of gravity signal in the two adjacent profiles (L9 and L10). In addition, the sinuosity index value of 2.33 in this transition zone is much greater than the expected values for active faults. Finally, we observed lower–middle Pleistocene fluvial and alluvial sediments on the hanging wall of segment 2, whereas these deposits disappear southward in segment 3. In segment 3 the hanging wall deposits are middle–upper Pleistocene sediments. This suggests a larger uplift activity in segment 2, capable of preserving older coverage.
To the south, the separation of segments 3 and 4 is mainly argued on the longitude discontinuity of the fault scarps observed on the surface, and on the gravity-derived basement morphology. An example of this discontinuity is represented by the intrusion of the Miocene La Obra granite (Fig. 8). This more competent unit may be responsible for the offset in the rupture plane.
Based on the arguments listed previously, our first-order approximation states that segments 2, 3, and 4 behave as independent ruptures, where each one can generate a similar characteristic earthquake. In addition, the segmentation defined in this work is similar to a first-order approximation with the defined segmentation in a previous work using a topographic analysis (Rauld, 2011).
An important discussion is how independent the rupture of segments 2, 3, and 4, which are separated by less than 3 km, could be. In this scenario it has been suggested that the activation of one segment can trigger the displacement of the adjacent segment (Wesnousky, 2008). However, while this behavior is evident in 60 % of cases in strike slip faults, it is not necessarily applicable to reverse faults (Wesnousky, 2008). Some examples of continuous ruptures associated with a specific earthquake are Chi-Chi, Taiwan, 1999 (Chen et al., 2001); Marryat Creek, Australia, 1986 (Machette et al., 1993); Mikawa, Japan, 1945 (Wesnousky, 2008); and El Asnam, Algeria, 1980 (Yielding et al., 1981). In the last case segmented ruptures were formed, but were produced by several events.
In order to discuss the potential activation of several SRF segments, we propose possible scenarios. One possibility is the triggering of a segment given the displacement in the adjacent segment. This case does not imply more hazard because these events are not simultaneous. Another possibility is that at deeper levels the SRF behaves as a single unit, but its stress releases are discontinuous in space at the surface. One example of such a behavior in reverse faults was observed in the Tennant Creek earthquake in Australia. During this earthquake a single event generated a discontinuous rupture at the surface (Crone et al., 1989). According to Crone et al. (1989), this discontinuity was produced by the existence of an along-strike rupture barrier. Until now there is no geological evidence in the SRF of a rupture barrier like the one identified in the Tennant Creek case. In this regard, our results suggest that deformation is accommodated in several parallel faults that reach the surface. In fact, gravity profiles (Fig. 5) and TEM imaging (Fig. 3) demonstrate the presence of several parallel faults that cut the upper level of the basement, but in some gravity profiles some of these faults do not have an along-strike continuity, suggesting that all of these parallel ruptures stop in those places (see ”no slip” places in Figs. 6c and 9). In addition, the observed near-surface displacement has been estimated by means of the time-integration methodologies, gravity modeling, and sinuosity index going back at least 100 ky. These observations are not consistent with a continuous and homogeneous displacement along the whole fault trace in a time window (> 100 ky) that must involve the occurrence of several characteristic earthquakes.
Although we cannot rule out a single rupture of the whole SRF segment, our
evidence consistently favors the occurrence of a single-segment
characteristic earthquake with a rupture length of
The PGA modeling results are similar to the empirical PGA observed in other
reverse earthquakes. Examples of these are the Chūetsu, Japan,
The range of the PGA values modeled in this work, PGA > 0.3 g at
distances shorter than 10 km from the fault scarp, are similar to the
previous work done at the SRF (Pérez et al., 2014), up to 0.2 g in the nearby
10 km from the fault. Largest values are also similar, PGA
Despite the differences in the maximum earthquake,
Based on the characteristic earthquake definition, we present the PGA
response associated with the SRF in Fig. 10 (described in detail in Sect. 4.7). To determine the expected damage, we need appropriate fragility curves
for the study zone and the corresponding building typology, which are not
available. Alternatively, we have the chance to compare the expected
acceleration with the reported effects of the Maule 2010 earthquake
(
The extremely high risk zones of the area (hanging wall filled with sediments) are located to the east of segments 2 and 4 (Fig. 10). The hanging wall of segment 2 is almost completely urbanized now, mainly with houses of one or two floors, with better resistance than the buildings given their larger rigidity. The other extremely high risk zone on the hanging wall of segment 4 has few constructions and until 2015 was mostly a low-density urbanized zone. Given this scenario, a successful mitigation measure must limit the building construction in these areas or at least not allow unreinforced masonry buildings. In addition, the norm must prohibit building in the proximity of the surface rupture zone, with special emphasis on public buildings, like hospitals or schools, and industrial buildings that may cause major damage, such as gas stations or nuclear research plants.
Natural seismicity registered in a 1-year local network is compatible with SRF activity. However, stress release as seismic activity along the SRF is secondary compared to the activity observed in other sectors near Santiago.
Geophysical and geomorphological evidences suggest that the SRF is segmented
into four subfaults that are most likely activated independently. Under this
scenario a characteristic earthquake of magnitude
Based on the TEM imaging, Apoquindo hill outcrops, and seismic evidence, the SRF is a high-angle structure.
If the SRF is activated, it can produce building collapse; therefore, it is necessary to take preventative actions to avoid catastrophic damages. In particular, construction in the rupture zone must be highly restricted, and construction of unreinforced masonry buildings in hanging walls filled with sediments must be limited.
The integrated methodology applied in this study provides a valuable tool to estimate the seismic risk associated with crustal faults with low slip rate and subtle surface evidences.
The data sets “gravity database” and “seismic database” are available in the Supplement material.
We want to thank the important field support provided by A. Mella, A. Bosh, N. Moraga, G. Sielfeld, I. Santibañez, B. Perez, S. Pérez, R. Figueroa, and M. Lizama. T. García kindly provided field support as well as her expertise in geophysical software. We thank G. Cassasa and A. Yañez for providing their homes to install a seismic station for over a year, and ENERGÍA ANDINA for providing their seismometers. CG5 gravimeter was provided by CEGA, FONDAP-CONICYT project no. 15090013. TEM experiment was partially supported by Fondecyt project no. 1141139. DICTUC S. A and CIGIDEN (FONDAP-CONICYT project no. 15110017) provided economic support for the development of this work. We also thank R. Rauld and G. Vargas for improving the quality of the paper with their discussion and precise comments. Finally, we appreciate the discussions with G. Arancibia, J. Cembrano, T. Roquer, and all of the emerging Geosciences group at PUC.Edited by: B. D. Malamud Reviewed by: R. Rauld and G. Vargas Easton