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Contemporaryseismicity,1983-1993,anditsimplicationstoseismichazardintheCentralFrontRange,Colorado

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CONTEMPORARY SEISMICITY, 1983 - 1993, AND ITS IMPLICATIONS TO SEISMIC HAZARD IN THE CENTRAL

FRONT RANGE, COLORADO

Jacqueline D.J. Bott California Geological Survey, 165 Berry Street, Suite 210, San Francisco, CA 94107

Ivan G. Wong

Seismic Hazards Group, URS Corporation, 500 12th Street, Suite 200, Oakland, CA 94607

Jon Ake U.S. Bureau of Reclamation, Denver Federal Center, P.O. Box 25007, Denver, CO 80225

Key Terms: seismicity, focal mechanisms, potentially active faults,

tectonic stress, seismic hazard

ABSTRACT The occurrence of the 1882 M 6.6 earthquake suggests that there may be a significant, albeit poorly quantified, seismic hazard in the Front Range of Colorado. In an effort to better characterize this hazard in the central Front Range, we have used a joint hypocenter-velocity inversion to improve hypocentral locations and derive a velocity model from a well-recorded set of microearthquakes (Richter magnitude [ML] ≤ 3.3). The dataset consists of more than 1,100 events recorded by Microgeophysics Corporation (MGC) from 1983 to 1993 for the Denver Water Department. About 50 stations from the MGC network were used in this study. A 1D joint hypocenter-velocity inversion was first used to find the best estimate 1D velocity model for the region and the dataset. These results were then used as a starting model for a 3D inversion.

The majority of the more than 800 relocated earthquakes show no definitive association with the surface traces of mapped “potentially active” faults (Miocene and younger), although some events cluster at the ends of a few of these faults and sometimes along them. However, seismicity occurring in the vicinity of, and not apparently along, late Quaternary faults is the typical pattern observed throughout much of the interior western U.S. A significant aspect of the central Front Range seismicity is its persistent, moderate level of activity relative to many other portions of the Rocky Mountain region. Similar seismically active areas in the western U.S. are generally confined to areas of known late Quaternary faulting. In general, focal mechanisms are consistent with normal and strike-slip faulting in the western portion of the study area and reverse and strike-slip faulting in the eastern portion along the N- to NW-striking faults.

Although paleoseismic investigations in the past 20 years indicate that displacement along many of the faults (e.g., Floyd Hill, Oil Creek faults) has not occurred in the past 100,000 years or so, we believe their earthquake potential still remains unresolved given our observations of the contemporary seismicity and also the paucity of late Quaternary sediments in the region. Three possible explanations that would be consistent with the paleoseismic evidence are (1) deep-seated (> 15-18 km) rupture resulting in little to no surface displacement along the faults, (2)

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long recurrence intervals of several tens of thousands to more than 100,000 years, or (3) the Miocene and younger faults are not seismogenic and that earthquakes like 1882 occur on buried or yet undiscovered faults. INTRODUCTION Historically, central Colorado has experienced only low levels of naturally-occurring seismicity, with the 1882 earthquake of moment magnitude (M) 6.6 ± 0.6 being the most significant event to shake the Denver metropolitan area (Kirkham and Rogers, 1985, 1986; Spence et al., 1996) (Figure 1). This observation excludes the Rocky Mountain Arsenal (RMA) induced earthquakes of the 1960’s which began shortly after the injection of waste fluid into the crust (Healy et al., 1968; Major and Simon, 1968; Herrmann et al., 1981; Hsieh and Bredehoeft, 1981). Twelve events in the sequence of body-wave magnitude (mb) 4 to 5 generated significant ground shaking in Denver (Modified Mercalli [MM] intensity ≥ VI) (Kirkham and Rogers, 1985). With the exception of the 1882 event, the largest known earthquake to have occurred within the Front Range was a mb 4.0 earthquake that occurred on 25 December 1994 near Castle Rock at a depth of 23.5 km (Kirkham and Rogers, 2000). This earthquake shook an area of 1,700 km2 along the boundary between the Front Range and the Great Plains (see further discussion).

In contrast to the sparse historical seismicity, microearthquake monitoring by Microgeophysics Corporation (MGC) in the Front Range, west and southwest of Denver from 1983 to 1993 for the Denver Water Department (DWD), has revealed a surprising moderate level of seismicity (Richter magnitude [ML] ≤ 3.3). About 2,300 local events (defined as S-P time ≤ 3 sec) were detected during the 10-year period from which about 1,000 events were located by MGC (1994). In some instances, planar steeply-dipping clusters of seismicity appear to be associated with the surface traces and subsurface projections of Miocene and younger faults identified by Kirkham and Rogers (1981) as “potentially active” faults (Unruh et al., 1996). In the study area, the most prominent of these faults include the Golden, Floyd Hill, Kennedy Gulch, Ute Pass, Oil Creek and Rampart Range faults. These faults appear to be characterized by very low activity rates with long recurrence intervals of more than tens of thousands of years (Widmann et al., 1998).

Unruh et al. (1998) suggest that the slow extensional deformation in the southern Rocky Mountains is driven by lithospheric gravitational potential energy and is accommodated locally by reactivation of pre-existing crustal faults. Although the southern Rocky Mountains are characterized by high potential energy, Jones et al. (1996) suggested that the region is deforming relatively slowly because of the strong underlying lithosphere.

The study of intraplate microseismicity and its relation to the reactivation of older faults is important for the Front Range because of the implications to seismic hazard. The 1882 earthquake is now thought to have occurred somewhere along the northern Front Range (Kirkham and Rogers, 1985, 1986; Spence et al., 1996), and thus the potential for similar-sized events may exist elsewhere along the Front Range. If we can identify which faults might be reactivated, even at the microearthquake level, our understanding of any potential earthquake threat to the rapidly growing central Front Range region can be improved.

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Figure 1. Historical seismicity (1870-2003) in Colorado. Location accuracy varies greatly over the state with large uncertainties (> 50 km) possible for pre-instrumental (pre-1960’s) earthquakes.

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Thus, the objectives of this study were to (1) improve the locations of earthquakes recorded by MGC using a joint-hypocenter velocity inversion, (2) better characterize the local crustal velocity structure, and (3) compute reliable focal mechanisms. Based on the improved hypocentral locations, possible associations between the earthquakes and the mapped Miocene and younger faults will be evaluated. The style of faulting determined from focal mechanisms can help characterize the regional state of stress in this transition zone and thus help in identifying which faults might be preferentially reactivated. Based on these results, the seismotectonic setting of the central Front Range and hopefully the seismic hazard for the region can be better characterized. The following paper summarizes the approach, the results of the analyses, and our interpretations of those results. A more complete description of this study, which describes in better detail the inversion for crustal velocity structure, is contained in Bott et al. (2003). APPROACH Although the number of MGC stations and the network configuration varied throughout the 1983 to 1993 monitoring period (MGC, 1994), in general, the network covered the portion of the Front Range between Boulder in the north and Colorado Springs to the south and from Lakewood westward for a distance of about 60 km (Figure 2). More than 50 station locations were occupied over the 10-year period, but the station configuration changed with only about 20 to 30 stations operational at any one time.

Only limited information gathered by MGC from the Front Range network operations has been published (Butler and Nicholl, 1986). However, the whole Front Range data set has been made available recently by DWD for use in seismic hazard studies of other facilities within central Colorado (e.g., Unruh et al., 1996; Bott et al., 1996) and this has renewed interest in this important data set.

Joint Hypocenter-Velocity Inversion

In the routine processing of earthquake locations, MGC used the program HYPO71 (Lee and Lahr, 1972), and a velocity model consisting of a 1.95 km-thick layer with a P-wave velocity (VP) of 5.1 km/sec over a half-space velocity of 5.95 km/sec. The velocity model was originally derived from blast data. Arrival-time data from a different set of blasts were used to verify and constrain the velocities in the uppermost layers of the crust. In this study, a 1D joint hypocenter- velocity inversion was first performed to determine the minimum VP model (FR1D), using the program VELEST (Kissling, 1988) and the 190 best-located events. Velocity layers were fixed in the near surface down to a depth of 4 km based on the blast results. Layer thickness varied from 1 to 2 km thick in the upper crust increasing to 4 to 6 km thick in the mid-crust. The best estimate 1D minimum model obtained was similar to the models of MGC and Prodehl and Lipman (1989) for the upper 10 km of the crust. However, slightly higher VP values (6.15 to 6.25 km/sec) were found between the depths of 10 and 26 km. No rays traversed layers below a depth of 26 km and so the velocities at these depths were not constrained. Based on the arrival time data, the VP/VS ratio is 1.71 (shear-wave velocity VS), lower than the 1.87 originally used by MGC. The VS was calculated assuming this ratio for the upper crust.

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Figure 2. Front Range stations operating for more than 1 year and 12-km arc around each station within which all ML 1.0 events are thought to have been detected. Miocene and younger faults are from Kirkham and Rogers (1981).

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The FR1D model incorporates station corrections, which were compared with the surface geology. Most stations located on the 1.08 billion-year-old Pikes Peak batholith (Smith et al., 1999) have the highest positive station corrections, implying slower velocities than the FR1D model predicts. In contrast, those stations located on older Precambrian metamorphic rocks (1.4 to 1.8 billion-years-old) have negative station corrections, implying faster velocities than FR1D predicts. MGC also used station corrections, which generally reflect the difference in station elevation and the elevation datum (2.5 km above mean sea level) since topography is not accounted for in HYPO71. Station elevations are accounted for in FR1D, and so the calculated station corrections should only account for the difference in the underlying rock velocities. Earthquake relocations of the best-located 190 events used in the 1D inversion moved on the order of 2 to 3 km from their original locations and a tightening of clusters of events was observed. The velocity model was fixed and the inversion repeated with all station elevations set to zero, to produce a set of station delays that could be used with program HYPO71 (velocity model referred to as FR1D-H71). Relocations of 11 known blasts using FR1D-H71 showed improvement on average of 0.9 km over locations using the MGC velocity model when compared to their known locations.

All events thought to be earthquakes were relocated with the new velocity model (FR1D-H71). An improvement was observed in the locations based on the average rms and location quality when compared to those computed using the MGC model. Earthquakes in the northern part of the network generally occur at shallow depths (< 15 km) and are diffusely distributed, whereas earthquakes in the south mostly occurred at depths greater than 8 km and sometimes in clusters or as swarms. In the north, the geology is comprised of Precambrian metamorphic rocks, whereas to the south, the network is underlain by the younger anorogenic granite of the Pikes Peak batholith. Structurally, the metamorphic rocks under the northern part of the network are cross-cut by numerous NW-SE-striking faults, whereas the batholith is cut by fewer faults and these generally strike N-S or are ring fractures associated with the emplacement of plutons within the batholith.

The 190 relocated earthquakes and the FR1D velocity model were used as initial input for a 3D joint hypocenter-velocity inversion using the program FDTOMO (Benz et al., 1996). Crustal velocities were determined for the middle and upper crust, down to about 26 km in some areas. Moho refractions are not found at the epicentral distances (less than 100 km) used in this data set. The relocated hypocenters were not too different from the locations using FR1D but are generally deeper, with average shifts of 1.0 km horizontally and 1.4 km vertically. Relocations of the same set of known blasts showed significant improvement with the average location being within 0.5 km of the actual site. Final relocations of more than 800 earthquakes using FR3D are shown in Figure 3.

The tomographic velocity model provides an interesting image beneath the Front Range, with a region of slightly lower velocity directly under the outcrop of the Pikes Peak batholith. This lower velocity area may define the shape of the batholith at depth. This region also correlates well with the area for which positive station corrections were determined using FR1D. It is observed that the microseismicity in the vicinity of the Pikes Peak batholith is restricted to clustering of events around and below the low-velocity region. Significant deformation may not be occurring internally within the batholith but may be localized along distinct fault zones around its edges and below it.

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Figure 3. Relocations of all tectonic earthquakes using FR3D showing shallow (≤ 8 km) events in blue and deep events (> 8 km) in red. Dotted line is the boundary of the Pikes Peak batholith. Unnamed faults not labeled. Cross-sections are shown in Figure 5. The location of the 25 December 1994 mb 4.0 Castle Rock earthquake is also shown.

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In the northern part of the network, the microseismicity is diffusely distributed both in space (Figure 3) and time. The earthquakes could be occurring on the many small faults that crosscut the metamorphic rocks in this region. Thus, the northern and southern regions appear to be deforming differently possibly due to their differing geology, and the location and orientation of pre-existing faults, despite the fact that both regions probably are being subjected to the same extensional tectonic stress field. An area of slightly higher velocity lies directly below the Pikes Peak batholith (> 6 km), which we believe is defined by the slightly lower velocities as described earlier. We speculate that this pattern could be revealing the existence of deep thrust structures under the Front Range, or a mafic igneous body. The latter is consistent with the model for the genesis of the A-type granites suggested for the Pikes Peak batholith (Smith et al., 1999) and is our preferred model. This hypothesis is also supported by the gravity data as other authors have suggested that the slight gravity high above the batholith indicates that the normally low-density granite (and thus an expected negative gravity anomaly) is somehow compensated for at depth by a denser rock mass.

Focal Mechanisms

An objective in this study was to use the improved locations to better determine single-event focal mechanisms. Focal mechanisms were computed for the best-located events (A and B quality locations) using the 3D velocity model and the program FPFIT (Reasenberg and Oppenheimer, 1985). Only mechanisms with at least 10 first motions were calculated.

A known problem with the data set was that several stations were believed to have polarities that were reversed from the standard assumption that an “up” motion was compressional and a “down” motion was dilatational. Unfortunately, the dates of station reversals were not completely documented and so we attempted to reconstruct the network station polarity history using the available data. To further interpret this data set, it was necessary to select a subset of the better-constrained events and stable solutions. To accomplish this, events with focal depths less than 5 km were removed as these depths are not as well constrained and thus probably not as accurate as those of the deeper events, based on station spacing.

Due to the difficulty of assessing the station polarity history by sifting through the thousands of original smoked paper records to identify known blasts, another approach was employed to utilize the available data. Stations recording the known and suspected blast events that did not have compressional first motions were noted. A list of suspect stations and their estimated time periods of reversal was compiled and used as input to the FPFIT program for a second focal mechanism determination. The original smoked paper records were checked for a few selected months to verify some of the station reversals.

The FPFIT program was rerun using the 3D-velocity model incorporating the proposed station reversal history and also using the 1D and the original MGC models. Focal mechanisms for events which appeared relatively stable despite a change in velocity model, and the incorporation of the station reversals, were selected to represent the final subset of data.

Focal mechanisms were also calculated for several events for a range of fixed focal depths to evaluate their sensitivity to depth. Several events, chosen to represent a range of faulting style

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and focal depth, were relocated with the focal depth fixed in 1-km increments from 5 to 25 km. The 1D velocity model was used for these analyses due to speed of running HYPO71. It was noted that for each of these events, their rms error occurred in a pronounced minimum in the rms error versus depth curve indicating that the actual focal depth was stable. The focal mechanisms were also found to be relatively stable for most of the events over the entire depth range, except in a few cases where the mechanism appeared to be affected for depths close to a velocity interface. The final filtered data set consists of 25 focal mechanisms: 12 normal, 7 reverse, 4 strike-slip, and 2 oblique-slip solutions (Figure 4). RESULTS AND DISCUSSION The earthquake relocations using the 3D model differ from the starting MGC locations, but mostly in their depth. Based on the relocation of the 11 quarry blasts, we believe the epicentral locations are accurate to within about 0.5 km, at least in the central portion of the network. The location accuracy probably deteriorates toward the edge and outside of the region covered by the network. However, it is noted that the locations of earthquakes occurring in the vicinity of the RMA, which are outside the network coverage, appear to be much closer to the fault than the original MGC locations (Figure 3), which indicates an improvement using the 3D velocity model. The depths of the RMA events are probably not well resolved and a change in the velocity with depth could have a large effect. In our model, the velocity increases with depth, whereas the MGC model has a constant half-space velocity below the top 2 km of the crust. It is therefore not surprising that the relocated earthquakes are much deeper than the original MGC locations. We believe the relocated events within the bounds of the network have a focal depth accuracy of ± 2 km.

Distribution of Seismicity

Based on the final relocations, microearthquakes appear to be distributed across the region covered by the network (Figure 3). A large concentration of events occurs within a 60-km-long NNW-SSW-trending rectangular area bounded partially on the west by the Floyd Hill and Kennedy Gulch faults, and on the east by the Golden and Jarre Creek faults (Figure 3). The 1994 Castle Rock earthquake, which was located by MGC (Kirkham and Rogers, 2000), appears to have occurred along the eastern margin of the Pikes Peak batholith (Figure 3). On the eastern edge of the batholith, along the Rampart Range fault, a small cluster of shallow events is observed orientated NW-SE rather than N-S parallel to the fault (Figure 3). This lineation could be an artifact of the station locations since this cluster is located towards the edge of the network region or it could be in fact a NW-SE-striking unmapped fault or fault zone.

The depth distributions of the relocated earthquakes are shown in Figure 5. Events are shown within 5 km of the cross-section. The majority of earthquakes occur at depths of less than 15 km. Cross-section C-C' shows an interesting and significant increase in the maximum depth of events from west to east across the network. The deepest earthquakes occur at depths down to 27 km beneath the northern edge of the Pikes Peak batholith (Figures 3 and 5, Sections C-C' and F-F'). In the area underlain by the batholith, shallow earthquakes are fewer in number (Figure 5, Section D-D').

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1 8 15 22 2 9 16 23 3 10 17 24 4 11 18 25 5 12 19 6 13 20 7 14 21

Figure 4. Focal mechanisms calculated in this study. Compressional and dilatational first motions indicated by plus symbols and circles, respectively. Orientations of pressure (P) and tension (T) axes are also shown.

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Figure 5. Cross-sections showing earthquake locations from the 3D velocity inversion. Locations of cross-sections and fault abbreviations shown on Figure 3. PPB in Section F-F' indicates northern edge of Pikes Peak batholith.

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A NNW-SSE-trending cross-section parallel to the strike of the Front Range shows the same interesting distribution of earthquakes with depth (Figure 5, Section F-F'). In the north from a distance of about 0 to 70 km, the earthquakes are distributed somewhat uniformly throughout the top 15 km of the crust. Near the northern edge of the Pikes Peak batholith, a sudden increase in the depth of events is observed with most earthquakes occurring at depths greater than about 8 km to the south. The earthquakes then shallow towards the southern end of the network. Two distinct clusters are observed in Section F-F' (Figure 5) in the southern half of the network: one deep concentration just north of the batholith and one very shallow cluster at an along-strike distance of 100 to 120 km.

Earthquakes in the interior intraplate portions of the western U.S. are generally confined to the top 15 to 20 km of the brittle crust (Smith and Bruhn, 1984; Wong and Chapman, 1990). The observed unusually deep events (> 20 km) indicate that a relatively cold thick seismogenic crust may exist beneath portions of the central Front Range.

Relationship to Faulting and Focal Mechanisms

In general, there is no clear causal association between seismicity and known Miocene and younger faulting, although there may be a spatial coincidence in a few possible areas (Figure 3). Some relocated earthquakes appear to cluster at the northern end of the Golden fault, in the center of and at the southern end of the Floyd Hill fault, possibly along the southern portion of the Kennedy Gulch fault, at the northern end of the Ute Pass fault, just west of the center of the west-dipping Rampart Range, and at the southern end of the northeast-dipping Chase Gulch fault (Figure 3). This tendency for earthquakes to cluster at the ends of generally seismically quiescent faults has been observed in western Nevada (VanWormer and Ryall, 1980) and is consistent with the suggestion that stress accumulates at fault tips.

The apparent spatial coincidence of the epicenters with the ends of faults noted above, however, is not readily visible in cross-section, although the dips of these faults are poorly known. Possible exceptions are the northern end of the west-dipping Golden faults, the east-dipping Floyd Hill fault, the west-dipping Rampart Range fault, and possibly the east-dipping Chase Gulch fault (Figure 5).

Due to the distribution of stations, the focal mechanisms calculated in this study were confined to earthquakes located in the central portion of the network SSW of Lakewood (Figure 6). They exhibit a range of faulting styles including normal, reverse, strike-slip, and oblique slip (both normal and reverse) (Figure 4). Despite the variety of mechanisms, a pattern emerges that is not unexpected. West of the dashed line on Figure 6, the majority of focal mechanisms display normal faulting in response to NE-SW to E-W extension (minimum principal stress).

This tectonic stress field is consistent with the stress field inferred from other focal mechanisms in central Colorado (e.g., Wong, 1986; Bott and Wong, 1995). East of the dashed line, most of the focal mechanisms display predominantly reverse faulting in response to NW-SE to E-W-directed compression (maximum principal stress) as observed elsewhere in the Great Plains and Midcontinent (Zoback and Zoback, 1989). The principal stress directions for the four strike-slip mechanisms are consistent with a transition between the extensional stress field to the west and the compressional field to the east whereby the maximum principal stress rotates from a vertical

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Figure 6. Map of schematic focal mechanisms calculated in this study. Dashed line represents the possible boundary between the extensional Rocky Mountains and the compressional Midcontinent tectonic stress fields. Inward and outward arrows indicate maximum and minimum principal stress directions, respectively. Shaded areas in focal mechanisms are compressional quadrants.

orientation to a horizontal position (Figure 6). The focal mechanism for the 1981 Conifer earthquake (Butler and Nicholl, 1986), which exhibits strike-slip faulting in response to an E-W maximum principal stress, is consistent with this transition. Normal faulting focal mechanisms along the RMA fault (Herrmann et al., 1981) are east of the dashed line shown in Figure 6 suggesting that the transition from extension to compression may not run north-south consistently through the Front Range. The overall tectonic stress pattern is consistent with the

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large-scale pattern suggested by Zoback and Zoback (1989). An interesting but unresolved aspect of the reverse focal mechanisms is they are for events that appear to be occurring deeper (> 15 km) than the normal/strike-slip faulting earthquakes (< 13 km) to the west.

Implications to Seismic Hazards

A desire of this study was to better define the potential seismic hazards in the central Front Range by evaluating whether the relocated contemporary seismicity could be correlated with any normal/strike-slip faulting of the mapped potentially active faults identified by Kirkham and Rogers (1985), and whether focal mechanisms were consistent with these structures being seismogenic. In this regard, the results of these analyses are inconclusive. The majority of relocated seismicity does not occur in the vicinity of mapped faults. However, some earthquakes are observed near or along several faults particularly at their mapped terminations.

In an apparent contrast, paleoseismic studies suggest that the potentially active faults in the central Front Range are either inactive or very low activity faults. Investigations conducted by and on the behalf of DWD of the Floyd Hill (Dickson et al., 1986), Kennedy Gulch (Yadon, 1986), Ken Caryl (Dickson and Paige, 1986), and Oil Creek faults (Friedman, 1986) have revealed no evidence for displacement along any of these structures in the past 80,000 to 100,000 years and in some cases, significantly longer. Circumstantial evidence for Quaternary slip is confined to the southern end of the Ute Pass fault (Dickson et al., 1986).

Trenching studies by Dickson (1986) on the Rampart Range fault suggest that the most recent displacement took place between about 30,000 to 50,000 to as much as 600,000 years ago. No Quaternary displacement has been observed along the Jarre Canyon fault (Dickson et al., 1986). The Golden fault has been the target of numerous paleoseismic investigations and the results have been ambiguous (Widmann et al., 1998). Widmann et al. (1998) have classified the Ute Pass and Rampart Range faults (last displacement prior to 30,000 to 50,000 years BP) as Quaternary in age. Unruh et al. (1998) estimate vertical separation rates on faults in the Front Range from displaced Quaternary strata ranging from 0.001 to 0.1 mm/yr.

It is important to note that the near-absence of small magnitude seismicity along many of the faults in the central Front Range is in itself, not sufficient to rule out that the faults are seismogenic since fault-related microseismicity is the exception rather than the norm in the interior of the western U.S. (e.g., Smith and Arabasz, 1991; Wong and Olig, 1998). The observation that some seismicity is occurring at the ends of some faults, as observed in the Basin and Range Province, remains intriguing. Although partially a result of the detailed microearthquake monitoring, another more general aspect of the central Front Range is the moderate level of seismicity. In the authors’ experience, such a level is observed only in regions of late-Quaternary faulting. Focal mechanisms are also consistent with the NW- to N-striking Miocene and younger faults observed in the Front Range being reactivated in normal, oblique-normal, and/or strike-slip displacement, in the extensional western portion of the region.

In summary, we believe, despite the sparse paleoseismic evidence, the Miocene and younger faults in the region may still be seismogenic. The faults could be very low activity faults with recurrence intervals on the order of several tens of thousands to greater than 100,000 years. This could explain the lack of evidence for surface displacement along most of the mapped faults in at least the past 100,000 years. Alternatively, as suggested by Spence et al. (1996), deep-seated

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rupture similar to the 24-km-deep 1985 M 5.5 Laramie Mountains earthquake may be a common mode of deformation in the Front Range. This is consistent with the presence of an approximately 25 km-thick seismogenic crust in portions of the central Front Range. An event like 1882 initiating rupture at a depth significantly greater than 15 km would likely result in no observable surface displacement accounting for the very sparse evidence for late-Quaternary activity on the Miocene and younger faults. A third possibility is that large earthquakes in the Front Range are occurring on buried or as yet undiscovered faults.

In light of both the microseismicity data and past paleoseismic studies, key questions with regards to seismic hazards in the central Front Range still remain unanswered: (1) what is the maximum sized earthquake that can occur; (2) what are the sources of such events; and (3) what are their recurrence intervals. Of primary relevance to addressing these questions is the 1882 M 6.6 earthquake. If the 1882 earthquake occurred in the northern Front Range as previously suggested, we know earthquakes of this magnitude, at a minimum, will likely occur in the future. Characterization of seismogenic structures whether they be any of the currently mapped Miocene and younger faults, unmapped or poorly studied faults, or buried faults remains our best hope for answering these questions. ACKNOWLEDGEMENTS This study was supported by the U.S. Geological Survey under the National Earthquake Hazards Reduction Program Award 99HQGR0028. Additional funding was provided by the British Geological Survey and URS Corporation. We would like to thank Susan Steele Weir of the DWD for archiving and allowing access to the original data, and her strong support for this project throughout. Our thanks to Dave Butler and MGC for their efforts over the years operating the Front Range network and for unselfishly providing data and assistance. We would like to thank Harley Benz of the USGS for the use of his inversion code and for his assistance, Antonio Villasenor (USGS) for making programs and GMT plotting scripts available, and Ann Sheehan and John Nicholl for providing data, information, and advice. Melinda Lee, Mark Dober, and Fumiko Goss assisted in the preparation of this paper. Critical reviews of the USGS report and this resulting paper were provided by David Butler, Susan Steele Weir, Pat Rogers, and Vince Matthews and we appreciate their efforts. REFERENCES Benz, H.M., Chouet, B.A., Dawson, P.B., Lahr, J.C., Page, R.A., and Hole, J.A., 1996, Three-dimensional P and S wave velocity structure of Redoubt Volcano, Alaska, Journal of Geophysical Research, v. 101, p. 8111-8128.

Bott, J.D.J. and Wong, I.G., 1995, The 1986 Crested Butte earthquake swarm and its implications for seismogeneisis in Colorado, Bulletin of the Seismological Society of America, v. 85, p. 1,495-1,500.

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Bott, J.D.J, Wong, I.G., Ake, J., 2003, Upper crustal velocities and earthquake relocations in the central Front Range, Colorado, from 1D and 3D joint hypocenter-velocity inversions, unpublished report prepared for U.S. Geological Survey, NEHRP Award 99HQGR0028.

Bott, J.D.J, Wong, I.G., Ake, J., Unruh, J., Nicholl, J. and Butler, D., 1996, Contemporary seismicity of the central Front Range, Colorado (abs.), Geological Society of America, Abstracts with Programs, v. 28, p. A-283

Butler, D. and Nicholl, J.J., 1986, The Conifer, Colorado earthquake of November 2, 1981, in W.P Rogers and R.M. Kirkham (eds.), Contributions to Colorado Seismicity and Tectonics – A 1986 update, Colorado Geological Survey Special Publication 28, p. 145-157.

Dickson, P.A., 1986, Investigation of the Rampart Range fault at the Air Force Academy trench site, Colorado Springs, Colorado, in W.P. Rogers and R.M. Kirkham, (eds.), Contributions to Colorado Seismicity and Tectonics- A 1986 update, Colorado Geological Survey Special Publication 28, p. 211-227.

Dickson, P.A., Kewer, R.P., and Wright, J.E., 1986, Regional fault study: Central Front Range, Colorado, in W.P. Rogers and R.M. Kirkham, (eds.), Contributions to Colorado Seismicity and Tectonics- A 1986 update, Colorado Geological Survey Special Publication 28, p. 172-185.

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