ABSTRACT

Previous theoretical work has suggested that porous sedimentary rocks subjected to suffi ciently high strain rates may not form deformation band damage zones (DBDZs). This hypothesis is evaluated by an examina-tion of faults within the porous Navajo Sand-stone at the Upheaval Dome impact structure in Canyonlands National Park, southeastern Utah, where high strain-rate conditions are known to have occurred. We found no evi-dence for DBDZ formation along the acces-sible fault planes at Upheaval Dome. Instead, a layer of pulverized quartz grains within the Navajo Sandstone occurs adjacent to the fault planes. Similar material has been ob-served in association with dynamic rupture in crystalline and sedimentary rocks along the San Andreas fault in southern California and in metamorphic rocks along the Bosman fault in South Africa. Measured grain sizes obtained from the pulverized material col-lected at Upheaval Dome are consistent with strain rates of ~1–3 × 103 s–1. Pulverized sedi-mentary rocks along the San Andreas fault imply strain rates of ~10–2 to 101 s–1. Strain rates along the faults at Upheaval Dome are well above the average values associated with intraplate tectonics but are consistent with, or faster than, seismic slip rates along faults such as the San Andreas fault. Our results support the hypothesized rate dependence of deformation in porous rocks.

INTRODUCTION

Several mechanisms for faulting that depend on rock type are well known in the literature. In low-porosity crystalline (or compact) rock such as granite, faulting occurs as a result of sliding on preexisting cracks (Segall and Pollard, 1983;

Martel and Pollard, 1989; Lockner et al., 1991; Lockner, 1995) or by failure of strain-softening damage zones of cracks resulting from shear deformation (Wawersik and Fairhurst, 1970; Wawersik and Brace, 1971; Lockner et al., 1991). In contrast, rocks of substantially higher porosity (>5%; e.g., Paterson and Wong, 2005), such as sandstones, fi rst generate deformation band damage zones (DBDZs), identifi ed as a network of strain-hardening deformation bands with reduced porosity and grain reorganiza-tion, that subsequently become faulted (Aydin and Johnson, 1978; Shipton and Cowie, 2001, 2003). Shear deformation of compact rocks may thus be characterized as crack-dominated defor-mation, whereas shearing of high-porosity rocks (e.g., Paterson and Wong, 2005) is initially com-paction dominated. The compaction-dominated sequence is well documented in Navajo Sand-stone (Aydin, 1978; Aydin and Johnson, 1978, 1983; Davis, 1999; Davis et al., 1999; Shipton and Cowie, 2001), a porous, quartz-rich, cross-bedded sandstone unit, and Entrada Sandstone (Aydin, 1978; Aydin and Johnson, 1978, 1983; Fossen and Hesthammer, 1997) on the Colorado Plateau, as well as in other units throughout the world (Aydin et al., 2006; Fossen et al., 2007).

Recent experimental and theoretical work sug-gests that deformation mechanisms in a porous rock may depend on strain rate (Fossum and Brannon , 2006). Using a combination of visco-plastic numerical models and experimental vali-dation (Kolsky bar tests), Fossum and Brannon (2006) found for Salem Limestone, under the conditions they evaluated, that the character-istic response time for compaction-dominated deformation to fully develop is approximately three times longer than for dilation-dominated deformation. According to Fossum and Brannon (2006), at low, quasistatic strain rates (i.e., ~10–5 s–1), either deformation mechanism can potentially occur in a porous rock. At higher, dynamic, strain rates, however (i.e., 102 s–1 for Salem Limestone), compaction-dominated defor-

mation may not have suffi cient time to develop, so that deformation of a porous rock would occur with cracking-dominated mechanisms only, as in a compact rock. This intriguing and novel re-sult suggests that a transition from compaction-dominated deformation to cracking-dominated deformation in other porous rock types, such as the Navajo and Entrada Sandstones, may also occur as a function of strain rate.

To evaluate the hypothesis of a rate depen-dence of deformation in porous sedimentary rocks subjected to high strain rates, we exam-ined macroscopic and microscopic data col-lected in the fi eld along faults inferred to have formed under these conditions. The Upheaval Dome impact structure was chosen for this in-vestigation because high strain-rate conditions are known to have been applied at this location during its formation, and faults associated with the impact event are well documented (Kriens et al., 1997, 1999; Kenkmann et al., 2005). Rock units affected by the impact and preserved within the perimeter of Upheaval Dome include Navajo Sandstone, which provides an ideal lithol ogy for comparing faults formed at low strain rates in continental tectonic settings, such as the Colorado Plateau, to those generated at high strain rates.

FAULTING IN POROUS SEDIMENTARY ROCKS

The processes leading to fault formation in porous sedimentary rocks have been exten-sively studied, particularly within the Colo-rado Plateau of the western United States (Aydin, 1978; Aydin and Johnson, 1978, 1983; Davis, 1999; Davis et al., 1999; Shipton and Cowie, 2001, 2003; Schultz and Balasko, 2003; Schultz and Siddharthan, 2005; Okubo and Schultz, 2005; Fossen et al., 2007). In porous granular rocks such as sandstones, deformation band damage zones are the pre-cursors to faulting in porous rocks in tectonic

For permission to copy, contact editing@geosociety.org© 2011 Geological Society of America

1161

GSA Bulletin; May/June 2011; v. 123; no. 5/6; p. 1161–1170; doi: 10.1130/B30087.1; 7 fi gures; 1 table.

†E-mail: wendy.key@amec.com

Fault formation in porous sedimentary rocks at high strain rates: First results from the Upheaval Dome impact structure, Utah, USA

Wendy R.O. Key1,2† and Richard A. Schultz1

1Geomechanics–Rock Fracture Group, Department of Geological Sciences and Engineering, MS 172, University of Nevada, Reno, Nevada 89557, USA2Now at AMEC Geomatrix, Rancho Cordova, California 95670, USA

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1162 Geological Society of America Bulletin, May/June 2011

settings such as the Colorado Plateau (Aydin and Johnson, 1978; Johnson, 1995; Okubo and Schultz, 2005). Deformation bands are identi-fi ed as tabular discontinuities within a porous granular rock that accommodate strain through compaction, dilation, shear, or a combination of these (Antonellini et al., 1994; Aydin et al., 2006; Fossen et al., 2007; Schultz and Fossen, 2008). Among the several kinematic varieties of defor mation bands, the most common in-volve both compaction (associated with poros-ity reduction and grain reorganization within the band) and cataclasis (cracking, rotation, and interlocking of grains within the band) (e.g., Aydin et al., 2006; Fossen et al., 2007).

Individual cataclastic deformation bands (Fig. 1A) are commonly millimeters in width and accommodate shear displacement on the order of several millimeters (Aydin, 1978; Aydin and Johnson, 1978). Further deformation within the host rock is accommodated by the forma-tion of additional parallel or subparallel defor-mation bands (Aydin and Johnson, 1978). The presence of two or more deformation bands defi nes a DBDZ (Aydin and Johnson, 1978). Deformation band damage zones characteristi-cally display a lenticular or wavy pattern when observed in the fi eld in map view (Aydin, 1978; Aydin and Johnson, 1978; Schultz and Balasko, 2003) and a backward-breaking linking geom-etry in cross section (Davis, 1999; Schultz and Balasko, 2003; Okubo and Schultz, 2005; Fig. 1B). Deformation band damage zones are typically on the order of decimeters to meters in total thickness (Aydin and Johnson, 1978; Ship-ton and Cowie, 2001). The total amount of off-set accommodated by a DBDZ can be estimated by summing the displacements of each individ-ual deformation band within the damage zone resulting in offsets up to several centimeters (Aydin and Johnson, 1978). To accommodate larger displacements on the order of decimeters to meters, widening of the DBDZ ceases and a slip surface forms; this surface eventually grows into a throughgoing fault that slices through the DBDZ (Aydin and Johnson, 1978, 1983; Shipton and Cowie, 2001, 2003; Schultz and Siddharthan, 2005; Fig. 1C).

PREVIOUS WORK AT UPHEAVAL DOME

Upheaval Dome is a circular structure ~5.5 km in diameter located on the Colorado Plateau in southeastern Utah (Fig. 2). It has been esti mated to be Jurassic (Alvarez et al., 1998) or Creta-ceous (Kenkmann et al., 2005) in age. Upheaval Dome is characterized by a complex central uplift, a ring syncline, and an outer circumfer-ential monocline dipping away from the center

of the structure (Shoemaker and Herkenhoff, 1983; Kriens et al., 1997, 1999; Jackson et al., 1998). A sedimentary sequence composed of the Permian White Rim Sandstone of the Cutler Group, Triassic Moenkopi and Chinle Forma-tions, and Jurassic Wingate Sandstone, Kayenta Formation, and Navajo Sandstone is exposed within the structure (Fig. 3). Radially oriented thrust faults with offsets ranging from 5 to 500 m are associated with structural thicken-ing within the central uplift (Kriens et al., 1999; Kenkmann et al., 2005). Concentric listric nor-mal faults are documented in the ring syncline (Kriens et al., 1999).

For decades, the origin of Upheaval Dome had been a source of debate; however, in recent years the theories for its formation have focused on salt intrusion or hypervelocity meteorite im-pact. According to the salt intrusion hypothesis, Upheaval Dome was formed by upward migra-tion of salt from the underlying Paradox For-mation, which has been regionally identifi ed as the source of numerous salt structures through-out the area (Harrison, 1927; McKnight, 1940; Mattox, 1975). Jackson et al. (1998) suggested that a salt diapir had passed through Upheaval

Dome’s center and had subsequently been pinched off and eroded away. Using structural and stratigraphic relationships, Jackson et al. (1998) suggested that the structure underwent prolonged growth beginning in the Jurassic and which continued for ~20 million years. They cited a concentric fold system composed of an outer rim monocline, a rim syncline, and an inner dome as geomorphic evidence for salt diapir-ism (Jackson et al., 1998). Additional evidence cited in favor of salt diapirism was synsedimen-tary deformation (i.e., truncations, channels, and growth folds) within the Chinle Formation, Wingate Sandstone, and Kayenta Formation, and emplacement of clastic dikes along faults in-ferred to have formed in the neck of the pinched-off salt structure (Jackson et al., 1998).

An origin by impact cratering was advo-cated by Shoemaker and Herkenhoff (1983) and Kriens et al. (1997, 1999) on the basis of detailed mapping of the structure and com-parisons with known terrestrial impact sites. Impact cratering is characterized by very high strain rates that generate extreme shock pres-sure and temperature conditions (Osinski and Spray, 2005; Pati and Reimold, 2007). Crater

CB

A

Figure 1. Deformation and faulting in porous Entrada Sandstone near the San Rafael Swell in southeastern Utah. (A) Several individual deformation bands; foot shown for scale. (B) A deformation band damage zone (DBDZ) with linking geometries visible in cross sec-tion; width ~1 m. Deformation bands in both (A) and (B) display positive relief in compari-son to the surrounding, undeformed Entrada Sandstone. (C) A normal fault with several meters of offset cutting the Entrada Sandstone above and the Carmel Formation below. The fault cuts through a DBDZ that is ~4 m in width (shown by arrow).

High strain-rate faulting in porous sedimentary rocks: First results from the Upheaval Dome impact structure

Geological Society of America Bulletin, May/June 2011 1163

formation can be divided into three sequential stages including contact and compression, exca-vation, and modifi cation (Melosh, 1989). Dur-ing the earliest contact and compression stage, kinetic energy is transferred from the impactor to the target rocks. Structures formed in the tar-get rocks during this high-pressure stage include shock metamorphism, cataclastic dikes, and shatter cones (French, 1998; Sagy et al., 2004). A bowl-shaped transient crater is formed during the excavation stage that is faulted during the modifi cation stage as a result of gravity-driven radial collapse (Croft, 1981a; Melosh, 1989). Structures forming during the modifi cation stage include concentric normal faulting, fold-ing, and uplift of the fl oor in the crater center (Croft, 1981b). Impact structures are known to form as a result of strain rates well in excess of 10–2 s–1 (Melosh, 1989; Pati and Reimold, 2007).

Huntoon (2000) presented a compilation of identifi ed strain indicators documented at Up-

heaval Dome to show that the structure’s mor-phology is consistent with proven impact craters rather than salt intrusion. He correlated the oc-currence of cataclastic dikes, shatter cones, thrust faults, and folds within the Wingate Sandstone with the excavation phase of crater formation. Evidence cited for the modifi cation phase in-cludes the central uplift, imbricated thrust sheets, outwardly plunging anticlines, ring syncline, and listric normal faults (Huntoon, 2000). Shocked quartz grains have also been identifi ed in inter-bedded sandstone units within the Kayenta Formation of Upheaval Dome’s ring syncline (Buchner and Kenkmann, 2008). Shocked quartz is considered to be an undisputed indicator of an impact origin (Ashworth and Schneider, 1985; Gratz et al., 1988; French, 1998).

Seismic-refl ection studies conducted across the structure have revealed that deformation decreases with depth and that layers above the Paradox Formation are relatively fl at lying,

favor ing an impact origin for Upheaval Dome and contradicting the salt intrusion model (Kanbur et al., 2000). Additional evidence sup-porting an impact origin for Upheaval Dome includes morphology and structural geology consistent with an impact structure (Shoemaker and Herkenhoff, 1983; Kriens et al., 1997, 1999), fault kinematics (Kenkmann et al., 2005), clastic dikes (Kenkmann, 2003), and spatially distributed cataclastic deformation bands in the Wingate Sandstone (Okubo and Schultz, 2007). Based on the evidence contained in the literature and summarized above, Upheaval Dome is best interpreted as an eroded complex impact crater, and faults mapped within the structure were formed in response to high strain rates generated during a hypervelocity impact event.

Much emphasis has been placed in terres-trial impact research on crater morphology (Shoemaker, 1960; Melosh and Gaffney, 1983; Melosh , 1989; Grieve, 1991; Morgan and Warner , 1999; Grieve and Therriault, 2004), mineralogical characteristics such as the pres-ence of high-temperature and high-pressure minerals and shock (or planar) deformation fea-tures in quartz and feldspars (Grieve and Ther-riault, 2004; Osinski and Spray, 2005; Buchner and Kenkmann, 2008), and presence of shatter cones (Grieve and Therriault, 2004; Sagy et al., 2004; Pati and Reimold, 2007). In contrast, studies of faulting at impact structures are less common. For example, Spray (1997) referred to some faults at impact craters as “superfaults,” or faults that accommodate large displacements (≥100 m) during single slip events in which velocities greater than 0.1 m/s occur. Detailed structural mapping studies have been conducted at several impact structures in sedimentary target rocks including Kentland (Laney and Van Schmus, 1978), Upheaval Dome (Kriens et al., 1997, 1999; Kenkmann et al., 2005), and Haughton crater in Canada (Osinski et al., 2005; Osinski and Spray, 2005). Structural and strati-graphic mapping at Upheaval Dome was con-ducted by Kriens et al. (1997, 1999), and in the central uplift by Kenkmann (2003), Kenkmann et al. (2005), and Scherler et al. (2006). Our paper builds on the geometry and kinematics of faults at Upheaval Dome by focusing on their potential strain-rate dependence.

OBSERVATIONS RELATED TO FAULTS AT UPHEAVAL DOME

Faults investigated in this study were iden-tifi ed at Upheaval Dome using the geologic map published by Kriens et al. (1999) with ac-cessibility evaluated initially by using Google Earth software and then verifi ed in the fi eld. Two faults, Fault 1 and Fault 2, were chosen for

UPHEAVALDOME

UTAH

TCPWRWR , TM,

JWJK

JN

Fault 2

Fault 1

1 km1 km

N

Figure 2. Oblique aerial view of the Upheaval Dome impact structure located in south-east Utah (see inset location map in the upper-left corner; photograph courtesy of Tom Till Photog raphy, Moab, Utah). The central uplift is characterized by the light-colored mounds in the center of the crater structure and is composed of the White Rim Sandstone (PWR) and the Moenkopi (TM) and Chinle (TC) Formations. The ring syncline is composed of the cliff-forming Wingate Sandstone (JW) located outside the central uplift, the topographic low formed by the Kayenta Formation (JK), and another cliff-forming unit, the Navajo Sand-stone (JN), located along the perimeter of the crater. The ring syncline extends outward from the Navajo Sandstone to the fl at-lying strata surrounding the impact structure. The loca-tions of the faults evaluated in this study, Fault 1 and Fault 2, are shown.

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1164 Geological Society of America Bulletin, May/June 2011

detailed study from a larger set of candidates that all cut Navajo Sandstone, given their ac-cessibility (i.e., not located along a cliff face) and state of preservation of fault-zone ma-terials (Fig. 2). Fault 2 has been mapped as a thrust fault (Kriens et al., 1999); however, the sense and magnitude of offset at Fault 1 are diffi cult to determine given the absence of pas-

sive markers . Faults observed in association with the impact at Upheaval Dome are known to have formed during a single event with no signifi cant offset along the faults generated subsequent to crater formation; these faults are thus distinct from those outside the structure that are not related to the impact (Kriens et al., 1999). Such conditions are ideal for the present

study, because the sequence of fault formation in porous Navajo Sandstone will be preserved with no addi tional deformation or overprinting.

Cataclastic deformation bands of post–Early Jurassic age have been documented at Upheaval Dome in the fi ne-grained, low-porosity Wingate Sandstone (Okubo and Schultz, 2007). Based on calculated stress magnitudes in excess of 0.7

0

100

200

300

400

500

600

700

800

900

Thickness(in meters)

PC

PWR

TM

TC

JW

JK

JN

JC

JE

Cutler Group- Red arkosic sandstone and conglomerate, sandy shale, and limestone.

Moenkopi Formation: Primarily red to brown shale, siltstone, and sandstone with gray lime-stone interfingering.

Chinle Formation: Partly silty and sandy variegated shales, with thin limestone and conglo-merate lenses.

Kayenta Formation: Red, gray,and purple shale andsandstone.

Navajo Sandstone: Porous,highly cross-bedded buff to redsandstone.

Wingate Sandstone: Massive, red, cliff-forming sandstone.

Carmel Formation: Red sand-stone, siltstone, and shale withbasal limestone. (Not drawn toscale.)

Entrada Sandstone: Light to red,usually massive sandstone. (Notdrawn to scale.)

Unit description

White Rim Sandstone observable at the top of the Cutler Group.

Per

mia

nTr

iass

icJu

rass

icA

ge

Uni

ts e

xpos

ed w

ithin

Uph

eava

l Dom

e

Deformation bands (Jurassic) documented in the J at Arches National Park.

E

Deformation bands (LateCretaceous to late Paleogene)documented in the J in western Utah.

N

Deformation bands (post–early Jurassic) documented in the J at Upheaval Dome.W

Additional comments

Figure 3. Stratigraphic column of units exposed at Upheaval Dome and younger eroded units after Jackson et al. (1998) (data from the Buck Mesa #1 well log [Fiero, 1958; Mattox, 1968; and Woodward-Clyde Consultants, 1983]). Measured thicknesses are reported beginning at the uncon-formity underlying the Permian rock sequence to the currently exposed surface. Unit descrip-tions are from Molenaar (1975). Ages of deformation are from Antonellini et al. (1994) (Entrada Sandstone), Davis et al. (1999) (Navajo Sandstone), and Okubo and Schultz (2007) (Wingate Sandstone).

High strain-rate faulting in porous sedimentary rocks: First results from the Upheaval Dome impact structure

Geological Society of America Bulletin, May/June 2011 1165

GPa, these deformation bands have been attrib-uted to the formation of Upheaval Dome during the early excavation stage of the impact event (Okubo and Schultz, 2007). During our inves-tigation, zones of deformation bands were also observed in the Navajo Sandstone in the ring syncline. The deformation bands were observed to be millimeters in width and occasionally ac-cumulated into small DBDZs on the order of several centimeters in width. Observed defor-mation bands and DBDZs within the Navajo Sandstone exhibit no evidence (i.e., slip sur-faces) for faulting.

The architecture of the faults we examined in Navajo Sandstone at Upheaval Dome is dis-tinctly different from that commonly observed in this unit elsewhere on the Colorado Plateau. Instead of DBDZs that were subsequently faulted, we observed a friable rock with a pow-dery texture exposed adjacent to distinct fault planes. At Fault 1 (Figs. 4A and 4B), this rock is white in color and ~2–6 cm thick. At Fault 2 (Fig. 4C), the rock is tan in color and ~2–3 cm thick. Varying degrees of friability within the Navajo Sandstone along observed faults at Up-heaval Dome were noted in thin section, with the most friable rock observed at Fault 1 when compared to rock at Fault 2 and the undeformed Navajo Sandstone. At both faults, the Navajo Sandstone is observed above the fault with sheared Kayenta Formation below; no DBDZs are observed adjacent to either fault in the Na-vajo Sandstone as is commonly the case else-where on the Colorado Plateau in these units. The friable rock was identifi ed as deformed Navajo Sandstone based on fi eld relationships shown in Figure 4, in which it was consistently observed between the fault plane and relatively undeformed Navajo Sandstone. Thin-section observations were also used to identify this material as Navajo Sandstone (i.e., quartz rich, high porosity). Iron oxidation coating the quartz grains was noted in thin sections of the friable rock collected along Fault 2, and likely is re-sponsible for the tan color observed.

In thin section the friable rock differs from fault gouge (Engelder, 1974) by the lack of signifi cant rotation or shear of the grains after fragmentation and by the absence of fabrics that would be indicative of shear. As shown in Fig-ure 5, the grains of the friable rock have clearly been shattered yet maintain their original grain boundaries. Shattered clasts within these grains have been visibly reduced to the microscale. Gouge zones are typically characterized by the progressive development of shear fabrics includ-ing en echelon R

1-shears, conjugate R

2-shears,

antithetic P- and X-shears, and Y-shears parallel to country rock interfaces (Gu and Wong, 1994). Fault material observed along Fault 1 and Fault 2

displayed no evidence for shear localization or cataclastic fl ow after fragmentation, and is there-fore not consistent with fault gouge.

We interpret the friable rock observed along fault planes at Fault 1 and Fault 2 as “pulver-ized” Navajo Sandstone, a material consistent with cracking-dominated deformation. Pulver-ized rock is defi ned by Dor et al. (2006) and Dor (2007) as rock that typically yields a white, powdery texture in the fi eld and is observed in thin section to be shattered in-place to the micro scale, while maintaining its original grain fabric with little to no evidence for rotation or shear after fragmentation. Five damage classes based on the intensities of fractured and pulver-ized rock have been identifi ed (Dor et al., 2006; Dor, 2007) and are summarized in Table 1. Weak and/or selective pulverization, in which some grains remain intact while others shatter, is identifi ed in thin section along the faults at Upheaval Dome, which corresponds approxi-mately to damage class IV as defi ned by Dor et al. (2006) and Dor (2007).

Pulverized rock along a fault is now recog-nized as distinct from gouge (Brune, 2001; Dor et al., 2006, 2009; Dor, 2007) and is considered to be diagnostic of high-velocity slip events along a fault (Brune, 2001; Wilson et al., 2005; Reches and Dewers, 2005; Dor et al., 2006, 2009; Dor, 2007). Damage along a fault plane has been shown to be asymmetric, with increased damage on the stiffer side of the fault (Dor et al., 2006, 2009). This is consistent with our observations at Upheaval Dome in which we fi nd rock pul-verization in the stiffer Navajo Sandstone above the fault plane and shear deformation in the softer clay and siltstone units of the Kayenta Formation below the fault plane.

Current research on high strain-rate faulting has linked dynamic slip events to pulverization in a variety of rock types ranging from granite to porous sedimentary rocks such as those found in the Juniper Hills and the Hungry Valley Forma-tions along the San Andreas fault (Brune, 2001; Dor et al., 2006, 2009; Dor, 2007), to quartzites along the Bosman fault in South Africa (Reches and Dewers, 2005; Wilson et al., 2005). The origin of pulverized rock has been closely asso-ciated with seismogenic faulting (i.e., dynamic slip events) (Prakash et al., 2008). Two mecha-nisms have been proposed for the formation of pulverized rock including a localized reduction in fault-normal stress during rupture passage (Brune, 2001; Wilson et al., 2005; Dor et al., 2006) or pulverization due to changing stress conditions localized at the rupture tip during earthquake propagation (Reches and Dewers, 2005; Wilson et al., 2005). Field studies by sev-eral research groups worldwide are demonstrat-ing that material along some major faults may

be pulverized rock, instead of gouge as previ-ously mapped or inferred (Brune, 2001; Dor et al., 2006), with important revised implica-tions for the mechanical behavior of faults.

ANALYSIS AND DISCUSSION

Grain-Size Analysis

Several different methods have been em-ployed to evaluate the reduced grain sizes of fault rocks including water disaggregation (An and Sammis, 1994; Reches and Dewers, 2005), sieve analysis (An and Sammis, 1994), Coulter-counter analysis (An and Sammis, 1994), and digital analysis of grain perimeter ratios from thin-section images (Dor, 2007; Dor et al., 2009). For this study, we chose to measure grain sizes directly from images taken from thin sec-tions. This method has been used in the evalua-tion of grain-size reduction within deformation and compaction bands (Sternlof et al., 2005; Torabi et al., 2007; see also Holcomb et al., 2007) and displays the grain sizes and packing conditions as they exist in the fi eld.

For this study, grain sizes were measured by summing the measured length and width of each individual grain and dividing the resulting value by two (Key, 2009). Pulverized grain sizes were measured petrographically from each of the indi vidual rock fragments observed within the original grain boundaries (see Fig. 5B). Grain sizes ranging from 3.75 × 10–6 m to 1.40 × 10–4 m were obtained for pulverized rock at Upheaval Dome. In order to compare the pulverized grain sizes to the original grain sizes prior to deforma-tion, measurements of the original quartz grains were also collected. These measurements were possible because the rock’s original fabric was maintained after pulverization, thus preserving original grain boundaries. Original grain sizes of the pulverized material were measured using the method described above.

Figure 6 presents a log-log plot of measured grain sizes of pulverized rock at Fault 1 (see Fig. 5). We obtain a two-dimensional (2D) power-law slope of 0.77 for the original grain sizes and 1.55 for the pulverized grain sizes col-lected from the Upheaval Dome samples. Al-though the regression lines shown in Figure 6 have a negative slope, the power law is reported as a positive value by convention to allow for more direct comparisons to previous work eval-uating the grain-size distribution of fault gouge (An and Sammis, 1994; Sammis and Steacy, 1995; Sammis and King, 2007). The 1.55 power-law slope obtained from the pulverized grain sizes is remarkably close to the 1.6 2D slope obtained from fault gouge reported by Sammis et al. (1987) and Sammis and Steacy (1995).

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1166 Geological Society of America Bulletin, May/June 2011

Undeformed Navajo Sandstone

Pulverized Navajo Sandstone

ShearedKayenta Formation

Fault Plane

B

C

Navajo

Kayenta

Undeformed Navajo

Sandstone

ShearedKayenta Formation

Pulverized Navajo Sandstone

A

Fault Plane

Figure 4. (A) Schematic diagram of Fault 1 with pulverized Navajo Sandstone shown in black just above the fault plane; fi eld notebook for scale. (B) Photograph of Fault 1 oriented facing south and away from the center of the crater (coordinates: 38°25′36.684″, 109°56′13.258″, North American Datum of 1983). Pulverized Navajo Sandstone is shown as the thin white unit (2–6 cm in width) to the right of the fi eld notebook and above the fault. In both (A) and (B), sheared siltstone of the Kayenta Formation is observed beneath the fault plane and to the left of the notebook. The Kayenta Formation–Navajo Sandstone contact left of the main fault plane is shown by the dashed line on the left of the cave. (C) Outcrop exposure of Fault 2; geologist, circled, is sitting on the fault plane separating Navajo above from Kayenta below.

High strain-rate faulting in porous sedimentary rocks: First results from the Upheaval Dome impact structure

Geological Society of America Bulletin, May/June 2011 1167

It has been suggested that the distribution of grain sizes within fault gouge can be correlated to strain-rate conditions (Sammis et al., 1987; An and Sammis, 1994; Sammis and Steacy, 1995; Sammis and King, 2007). For example, a 2D power-law slope of 2.0 has been reported in association with higher strain rates, with a 2D power-law slope of 2.6 associated with “low-strain gouge” (Sammis and King, 2007). The pulverized rock collected at Upheaval Dome is inferred to have formed in a high strain-rate

environment, although the power-law slope associated with our pulverized grain sizes is comparable to the 1.6 2D power-law slope for “low-strain gouge and breccia” reported by Sammis et al. (1987) and Sammis and Steacy (1995). Our data suggest a degree of nonunique-ness in grain-size distributions of pulverized rock and fault gouge.

Strain Rates

Results from laboratory and controlled blast-ing experiments used to investigate dynamic fragmentation in rock have defi ned a relation-ship between grain size and the strain rate re-sponsible for grain-size reduction (Grady and Kipp, 1987; Zhou et al., 2005). Blasting ex-periments were conducted within oil shales in Rifl e, Colorado, in order to prepare the rock for retorting (heating), allowing access to liquid oil (Grady and Kipp, 1987). Strain rates applied to the rock during these experiments ranged from 100 to 104 s–1 (Grady and Kipp, 1987). The rela-tionship derived by Grady and Kipp (1987) was most recently applied by Reches and Dewers (2005) in their study of dynamic rupture along faults. This relationship is given by:

d =20KIC( )ρCdε

⎡

⎣⎢⎢

⎤

⎦⎥⎥

2

3

, (1)

where d is the diameter of the fragmented grains, KIC is the mode-I fracture toughness of the unfractured grain, ρ is the rock density, Cd is the compressive (P-wave) velocity of Navajo Sandstone, and ε is the strain rate applied to the rock. The reduction of grain size during dynamic fragmentation has been found to be dependent on the kinetic energy applied to the rock affected. Therefore, the role of initial grain size is relatively insignifi cant in compari-son to the rock’s properties and the strain rate applied to the rock (Grady and Kipp, 1987; Reches and Dewers, 2005). Additionally, over-burden pressure is negligible in comparison to the extreme loading conditions applied to the

rock during dynamic fragmentation (Grady and Kipp, 1987).

Zhou et al. (2005) reevaluated Grady and Kipp’s (1987) relationship between grain size and strain rate by comparing data generated from one-dimensional (1D) armor ceramic sys-tem models to the data presented by Grady and Kipp. The comprehensive model presented by Zhou et al. (2005) took a numerical approach to the evaluation of dynamic fragmentation, considering a wide range of strain rates (101 to 5 × 106 s–1) and incorporating various spatial distributions of intrinsic defects and material strengths. Zhou et al. (2005) also considered elastic wave interaction with cohesive crack opening, resulting in a prediction of the num-ber of fragments a rock will break into under a given strain rate. Grain sizes calculated from these simulations, in which original grains with randomly oriented fl aws were subjected to high strain-rate conditions (defi ned as ≥103 s–1), re-vealed that grain sizes are reduced by ~40% compared to those predicted by Grady and Kipp (1987) for the same set of conditions (see Zhou et al., 2005). At low (quasistatic) strain rates (≤101 s–1) where only a portion of the strain energy goes toward the development of new cracks, fragment sizes are not dependent on strain rate, and a constant grain size is predicted (Zhou et al., 2005). We defi ne a parameter, f, which accounts for the shift in values obtained by Grady and Kipp’s relation (1987) as sug-gested by Zhou et al. (2005) to account for their more realistic (i.e., nonuniform) initial fl aw dis-tribution. Rearranged for strain rate, Grady and Kipp’s (1987) adjusted equation is given by:

ε =20KIC

ρCd

d

f

⎛⎝⎜

⎞⎠⎟

32

, (2)

where f = 1 in Grady and Kipp’s (1987) Equa-tion 1, and f = 0.4 following Zhou et al. (2005).

We have applied the adjusted Grady and Kipp (1987) relation (Equation 2) to the sam-ples of pulverized rock obtained from the fi eld

A0 μm 10 20 30

0 μm 10 20 30 B

Figure 5. Photomicrographs of quartz grains in undeformed Navajo Sandstone (A) and pulverized Navajo Sandstone (B). (A) Grains shown in the undeformed rock taken from outside the fault zone are subangular to subrounded and display little to no intra-granular fracturing. (B) Grains shown in the pulverized rock display signifi cant fractur-ing. Images taken in plane-polarized light.

TABLE 1. SUMMARY OF DAMAGE CLASSES RANGING FROM DEGREES OF FRAGMENTATION AND/OR FRACTURING (CLASSES I–III) TO DEGREES OF PULVERIZATION (CLASSES IV AND V) AS DEFINED BY DOR ET AL. (2006) FROM

OBSERVATIONS ALONG THE MOJAVE SECTION OF THE SAN ANDREAS FAULT IN SOUTHERN CALIFORNIA

Damage class DescriptionI Weak fracturing Rock with large macroscopic fractures exceeding background fracture density.II Fragmentation Rock that is fragmented with cm-scale fractures. Rock yields a rugged appearance with fragments that can be crushed

to smaller pieces by hand.III Intense fracturing Rock fractured at the grain scale. Rock yields a texture similar to grus and is characterized by rough, rounded surfaces.IV Weak and/or selective

pulverizationRock with grains fractured to the microscale. Some crystals or grains remain intact, some break along subcrystal or

grain fractures, and some yield a powdery texture. Rock is easily eroded with smooth and rounded outcrops.V Pervasive pulverization Rock with grains fractured to the microscale. All crystals or grains yield a powdery texture when crushed by hand.

Rock is easily eroded with smooth and rounded outcrops.Note: Damage classes are determined by the intensity of observed deformation in proximity to faulting and range from weak fracturing at the macroscopic scale to pervasive

pulverization at the microscopic scale (Dor et al., 2006; Dor, 2007).

Key and Schultz

1168 Geological Society of America Bulletin, May/June 2011

locations at Upheaval Dome in order to estimate strain rates along the faults at the time of im-pact. Our measured grain sizes are plotted in Figure 7 along with Equation 2. We chose val-ues of KIC of 10–100 MPa*m1/2, which brackets the value of 30 MPa*m1/2 used by Reches and Dewers (2005), ρ of 2250 kg/m3 appropriate to un deformed Navajo Sandstone (Sternlof et al., 2005), and Cd of 2483 m/s obtained using:

Cd = μ + λρ

, (3)

(Reches and Dewers, 2005). Values were ob-tained for Lamé’s coeffi cient (λ) and the shear modulus (μ) using:

λ = Eν1+ ν( ) 1− 2ν( ), (4)

μ = E

2 1+ ν( ) (5)

(Timoshenko and Goodier, 1970), with Young’s modulus (E) of 20 GPa and Poisson’s ratio (ν) of 0.2 (Sternlof et al., 2005) appropriate to Navajo Sandstone.

Using a representative grain size of 4 × 10–5 m for pulverized rock along faults at Upheaval Dome with f = 1 in Equation 2 (Grady and Kipp, 1987), we obtain an approximate strain rate of 88 s–1. Using maximum and minimum measured grain sizes of 3.75 × 10–6 to 1.40 × 10–4 m, strain rates ranging from 5 to 1 × 104 s–1 are obtained. Using f = 0.4, we obtain a strain rate of 2 × 101 s–1, with values ranging from 1 to 3 × 103 s–1 using maximum and minimum measured grain sizes.

To compare strain rates calculated along faults within Upheaval Dome to rates gener-ated elsewhere, strain rates associated with the pulverization of sedimentary rocks along the Mojave section of the San Andreas fault were also calculated. Doan and Gary (2008) esti-mated strain rates associated with pulverized rock along the San Andreas fault of at least 1.5 × 10–2 s–1 using high strain-rate laboratory testing methods. Using grain sizes ranging from 2.56 × 10–3 to 1.51 × 10–4 m obtained by Dor et al. (2009) for pulverized sedimentary rocks along the San Andreas fault and Grady and Kipp’s (1987) relationship (Equation 2, with f = 1), we obtain maximum and minimum strain rates ranging from 6 × 10–2 to 4 × 101 s–1. Using f = 0.4, we obtain strain rates ranging from 2 × 10–2 to 2 × 101 s–1. These values are consistent with those obtained by Doan and Gary (2008).

Our results suggest that strain rates calcu-lated for faults at Upheaval Dome (1 – 3 × 103 s–1) are consistent with, or faster than, strain

Pulverized grains

Original grains

1

10

100

1000

0.001 0.01 0.1 1

Grain size (d, mm)

Num

ber

(n)

y = 1.67x–0.77

y = 0.36x–1.55

10 –7

10 –6

10 –5

10 –4

10 –3

10 –2

10 –1

10 –15 10 –13 10 –11 10 –9 10 –7 10 –5 10 –3 10 –1 10 1 10 3 10 5

Gra

in s

ize

(d, m

)

Strain rate (s–1)

Upheaval Dome fault

KIc = 30 MPa*m1/2

San Andreas fault (Dor, 2007)

10

100

Figure 7. Plot of grain size versus predicted strain rate (Equation 2) compared to data col-lected along faults at Upheaval Dome (this work) and sedimentary rocks along the San Andreas fault (Dor et al., 2009). The dashed line on the right side of the plot represents KIC = 100 MPa*m1/2 for f = 1; the solid and dashed lines offset to the left represent KIC rang-ing from 10 MPa*m1/2 to 100 MPa*m1/2 for f = 0.4.

Figure 6. Log-log plot com par ing grain sizes in the undeformed and pulverized Navajo Sandstone as measured in thin section.

High strain-rate faulting in porous sedimentary rocks: First results from the Upheaval Dome impact structure

Geological Society of America Bulletin, May/June 2011 1169

rates calculated along the Mojave section of the San Andreas fault (2 × 10–2 to 2 × 101 s–1). Strain rates calculated for Upheaval Dome are up to 20 orders of magnitude larger than the high-est strain rates generated in intraplate tectonic settings (i.e., 10–20 to 10–17 s–1, Gordon, 1998) as well as those used in standard uniaxial and triaxial laboratory experiments (e.g., American Society for Testing and Materials [ASTM] stan-dard D7012, 2004). Our fi ndings are consistent with a general rate-dependence for deformation of porous sedimentary rocks with DBDZs de-veloping in response to strain rates <10–2 s–1 and pulverized rock forming in response to strain rates in excess of 10–2 s–1.

According to Wilson et al. (2005), rock pulverization is a function of strain rate and rupture velocity resulting from a single earth-quake event, but not a function of cumula-tive offset along a fault. Zones of pulverized rock are observed to approach 100–200 m in thickness along the Mojave section of the San Andreas fault (Brune, 2001; Dor et al., 2006, 2009), where cumulative fault offset is on the order of tens to hundreds of kilometers (Wilson et al., 2005). By contrast, zones of pulverized rock observed along the faults located within the ring syncline of Upheaval Dome are several centimeters in thickness. Maximum fault off-sets within the ring syncline at Upheaval Dome, where the faults evaluated in this study are lo-cated, are estimated to be on the order of meters to tens of meters (Kenkmann et al., 2005). Al-though preliminary, it appears that the width of pulverized zones might be related to the amount of fault offset.

CONCLUSIONS

Our fi ndings suggest that faulting in porous Navajo Sandstone within the ring syncline of Upheaval Dome was associated with the forma-tion of pulverized rock along fault planes and not by the formation of deformation band damage zones (DBDZs). This sequence of faulting con-trasts with that on the Colorado Plateau outside the impact structure, in which DBDZs form dur-ing the early stage of deformation. Using the re-lationships obtained by Grady and Kipp (1987) and refi ned by Zhou et al. (2005), we infer that the two faults studied at Upheaval Dome likely formed in response to strain rates ranging from 1 to 3 × 103 s–1. Using the same method, we obtain strain rates ranging from 6 × 10–2 to 2 × 101 s–1 for the Mojave section of the San Andreas fault that are consistent with dynamic earthquake slip rates. Our fi ndings support the inference of Fos-sum and Brannon (2006) that deformation mech-anisms in porous rock such as Navajo Sandstone are strain-rate dependent.

Pulverized rock has been documented in all rock types (i.e., sedimentary, crystalline, and metamorphic) where faulting is observed. Under-standing strain-rate conditions at an impact site provides a new source of information regarding the mechanical responses of rock types affected by deformation at high strain rates. Complex impact craters, such as Upheaval Dome, where faults display greater offsets at the center of the crater than near the crater rim, may provide ideal locations for testing the relationship of fault off-sets to widths of pulverized zones.

ACKNOWLEDGMENTS

We thank Thomas Kenkmann, Chris Okubo, and an anonymous referee for their helpful reviews, and Ken Herkenhoff, and especially Associate Editor W.U. Reimold , for comments that improved the clarity of the fi nal paper. Discussions with Ze’ev Reches, Yehuda Ben-Zion, John Louie, and Steve Martel were helpful during various stages of the work. This work was supported by a grant from National Aeronautics and Space Administration’s Planetary Geology and Geophysics Program.

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