Geological Criteria for Evaluating Seismicity Revisited: Forty Years of Paleoseismic Investigations and the Natural Record of Past Earthquakes (GSA Special Paper 479)

Geolooical Criteria for Evaluatin9 Seismicity Revisited

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Forty Years of Paleoseismic Investigations and the Natural Record of Past Earthquakes

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THE GEOLOGICAL SOCIETY OF AMERICA®

Special Paper 479

Edited by Franck A. Audemard M., Alessandro Maria Michetti, and James P. McCalpin

Geological Criteria for Evaluating Seismicity Revisited: Forty Years of Paleoseismic Investigations and the Natural Record of Past Earthquakes

edited by

Franck A. Audemard M. Fundación Venezolana de Investigaciones Sismológicas FUNVISIS El Llanito, Caracas 1073 Venezuela Alessandro Maria Michetti Dipartimento di Scienze Chimiche e Ambientali Università dell’Insubria Via Valleggio, 11 22100 Como Italy James P. McCalpin GEO-HAZ Consulting Box 837 Crestone, Colorado 81131 USA

Special Paper 479 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140, USA

2011

Copyright © 2011, The Geological Society of America (GSA), Inc. All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. In addition, an author has the right to use his or her article or a portion of the article in a thesis or dissertation without requesting permission from GSA, provided the bibliographic citation and the GSA copyright credit line are given on the appropriate pages. For permission to make photocopies of any item in this volume for other noncommercial, nonprofit purposes, contact The Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form for personal or corporate use, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA. GSA provides this and other forums for the presentation of diverse opinions and positions by scientists worldwide, regardless of their race, citizenship, gender, religion, sexual orientation, or political viewpoint. Opinions presented in this publication do not reflect official positions of the Society. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. GSA Books Science Editors: Marion E. Bickford and Donald I. Siegel Library of Congress Cataloging-in-Publication Data Geological criteria for evaluating seismicity revisited : forty years of paleoseismic investigations and the natural record of past earthquakes / edited by Franck A. Audemard M., Alessandro Maria Michetti, James P. McCalpin. p. cm. — (Special paper ; 479) Includes bibliographical references. ISBN 978-0-8137-2479-9 (pbk.) 1. Paleoseismology—Methodology. I. Audemard M., Franck A. II. Michetti, Alessandro M. III. McCalpin, James. QE539.2.P34G46 2011 551.22—dc22 2011007062 Cover: 2006 study at the northern strand of the Boconó fault at the Lagunillas pull-apart basin (state of Mérida, Mérida Andes, Venezuela). (Top) location of the trench site (trench under white tent), (bottom) trench view. Photos courtesy of Franck A. Audemard M.

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Contents

Introduction: Geological criteria for evaluating seismicity revisited: Forty years of paleoseismic investigations and the natural record of past earthquakes . . . . . . . . . . . . . . . . . . . . . . . . . 1 Franck A. Audemard M. and Alessandro Maria Michetti 1. Paleoseismicity of a low-slip-rate normal fault in the Rio Grande rift, USA: The Calabacillas fault, Albuquerque, New Mexico . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 J.P. McCalpin, J.B.J. Harrison, G.W. Berger, and H.C. Tobin 2. Late Quaternary earthquakes on the Hubbell Spring fault system, New Mexico, USA: Evidence for noncharacteristic ruptures of intrabasin faults in the Rio Grande rift . . . . . . . . . . 47 Susan S. Olig, Martha C. Eppes, Steven L. Forman, David W. Love, and Bruce D. Allen 3. Large-magnitude late Holocene seismic activity in the Pereira-Armenia region, Colombia . . . . 79 Claudia Patricia Lalinde P., Gloria Elena Toro, Andrés Velásquez, and Franck A. Audemard M. 4. Evidence of Holocene compression at Tuluá, along the western foothills of the Central Cordillera of Colombia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 91 Myriam C. López C. and Franck A. Audemard M. 5. Style and timing of late Quaternary faulting on the Lake Edgar fault, southwest Tasmania, Australia: Implications for hazard assessment in intracratonic areas . . . . . . . . . . . . . . . . . . . . . 109 Dan Clark, Matt Cupper, Mike Sandiford, and Kevin Kiernan 6. Multiple-trench investigations across the newly ruptured segment of the El Pilar fault in northeastern Venezuela after the 1997 Cariaco earthquake . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 Franck A. Audemard M. 7. Lake sediments as late Quaternary paleoseismic archives: Examples in the northwestern Alps and clues for earthquake-origin assessment of sedimentary disturbances . . . . . . . . . . . . . 159 Christian Beck 8. Late Pleistocene–early Holocene paleoseismicity deduced from lake sediment deformation and coeval landsliding in the Calchaquíes valleys, NW Argentina . . . . . . . . . . . . . . . . . . . . . . . 181 Reginald L. Hermanns and Samuel Niedermann 9. Rupture length and paleomagnitude estimates from point measurements of displacement— A model-based approach . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 195 Glenn Biasi, Ray J. Weldon II, and Kate Scharer

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The Geological Society of America Special Paper 479 2011

Geological criteria for evaluating seismicity revisited: Forty years of paleoseismic investigations and the natural record of past earthquakes Franck A. Audemard M.* Fundación Venezolana de Investigaciones Sismológicas (FUNVISIS), Final Prolongación Calle Mara, Quinta Funvisis, El Llanito, Caracas 1073, Venezuela, and School of Geology, Mines & Geophysics, Universidad Central de Venezuela, Ciudad Universitaria, Los Chaguaramos, Caracas 1010, Venezuela Alessandro Maria Michetti* Dipartimento di Scienze Chimiche e Ambientali, Università degli Studi dell’Insubria, Via Valleggio, 11, 22100 Como, Italy

ABSTRACT The identification of individual past earthquakes and their characterization in time and space, as well as in magnitude, can be approached in many different ways with a large variety of methods and techniques, using a wide spectrum of objects and features. We revise the stratigraphic and geomorphic evidence currently used in the study of paleoseismicity, after more than three decades since the work by Allen (1975), which was arguably the first critical overview in the field of earthquake geology. Natural objects or geomarkers suitable for paleoseismic analyses are essentially preserved in the sediments, and in a broader sense, in the geologic record. Therefore, the study of these features requires the involvement of geoscientists, but very frequently it is a multidisciplinary effort. The constructed environment and heritage, which typically are the focus of archaeoseismology and macroseismology, here are left aside. The geomarkers suitable to paleoseismic assessment can be grouped based on their physical relation to the earthquake’s causative fault. If directly associated with the fault surface rupture, these objects are known as direct or on-fault features (primary effects in the Environmental Seismic Intensity [ESI] 2007 scale). Conversely, those indicators not in direct contact with the fault plane are known as indirect or off-fault evidence (secondary effects in the ESI 2007 scale). This second class of evidence can be subdivided into three types or subclasses: type A, which encompasses seismically induced effects, including soft-sediment deformation (soil liquefaction, mud diapirism), mass movements (including slumps), broken (disturbed) speleothems, fallen precarious rocks, shattered basement rocks, and marks of degassing (pockmarks, mud volcanoes); type B, which consists of remobilized and redeposited sediments (turbidites,

*E-mails: [email protected], [email protected]; [email protected]. Audemard M., F.A., and Michetti, A.M., 2011, Geological criteria for evaluating seismicity revisited: Forty years of paleoseismic investigations and the natural record of past earthquakes, in Audemard M., F.A., Michetti, A.M., and McCalpin, J.P., eds., Geological Criteria for Evaluating Seismicity Revisited: Forty Years of Paleoseismic Investigations and the Natural Record of Past Earthquakes: Geological Society of America Special Paper 479, p. 1–21, doi:10.1130/2011.2479(00). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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Audemard M. and Michetti homogenites, and tsunamites) and transported rock fragments (erratic blocks); and type C, entailing regional markers of uplift or subsidence (such as reef tracts, microatolls, terrace risers, river channels, and in some cases progressive unconformities). The first subclass of objects (type A) is generated by seismic shaking. The second subclass (type B) relates either to water bodies set in motion by the earthquake (for the sediments and erratic blocks) or to earthquake shaking; in a general way, they all relate to wave propagation through different materials. The third subclass (type C) is mostly related to the tectonic deformation itself and can range from local (next to the causative fault) to regional scale. The natural exposure of the paleoseismic objects—which necessarily conditions the paleoseismic approach employed—is largely controlled by the geodynamic setting. For instance, oceanic subduction zones are mostly submarine, while collisional settings tend to occur in continental environments. Divergent and wrenching margins may occur anywhere, in any marine, transitional, or continental environment. Despite the fact that most past subduction earthquakes have to be assessed through indirect evidence, paleoseismic analyses of this category of events have made dramatic progress recently, owing to the increasingly catastrophic impact that they have on human society.

INTRODUCTION Paleoseismology is a currently growing and developing earth sciences discipline. The recognition that paleoseismology has in fact become a branch of learning in itself is a rather recent development (Slemmons, 1957; Allen, 1975; Sieh, 1978; Crone and Omdahl, 1987; Vittori et al., 1991; Yeats et al., 1997; Audemard, 1998; McCalpin, 1996, 2009). This is borne out by the recent publication of dedicated textbooks (Wallace, 1986; McCalpin, 1996; Yeats et al., 1997; Keller and Pinter, 1998; Burbank and Anderson, 2001; McCalpin, 2009) and special issues of scientific journals (e.g., Serva and Slemmons, 1995; Michetti and Hancock, 1997; Pavlides et al., 1999; Pantosti et al., 2003; Michetti et al., 2005a; Reicherter et al., 2009, this volume) or conference proceedings (e.g., Michetti, 1998; Cello and Tondi, 2000; HAN2000–PALEOSIS Project; Okumura et al., 2000, 2005; Toda and Okumura, 2010). Also, special issues or proceedings, and even books, on more specific paleoseismic issues have also come to light (e.g., Maltman [1994] on soft sediment deformation; Stratton Noller et al. [2000] on Quaternary dating techniques; Comerci [2001] on seismically induced mass movements; Gilli and Audra [2001] on speleoseismology; and Tappin [2007] on the sedimentary record of tsunamis and tsunamites). Finally, concerted efforts within the International Union for Quaternary Research (INQUA) community for several years have led to the recent proposal and acceptance (currently under validation) of a new macroseismic scale based exclusively on earthquake environmental effects, named the Environmental Seismic Intensity Scale 2007, known under the acronym ESI 2007 (Michetti et al., 2007). Most researchers agree that the main aim of paleoseismology is the identification and characterization of past earthquakes from geo-archives. In a broader sense, this type of research intends to establish the seismic history of a given fault or given region from deformed sediments or rocks, beyond the limited tempo-

ral resolution of instrumental and historical seismicity. However, this does not exclude major temporal overlapping with these two latter disciplines, as proposed by many authors (e.g., Gürpinar, 1989; Audemard, 1998, 2005; Levret, 2002; Carrillo et al., 2009). In regions of the world where historical and/or instrumental seismicity studies are really scarce, or where earthquake recurrence is too long to allow any detection of large past events, the study of past seismic behavior of faults has to rely exclusively on paleoseismology. Furthermore, paleoseismic research can even reveal very distinct trends of seismic activity over longer periods of time (longer than the current interglacial period), as is the case in large regions of the world subject to glaciation-deglaciation cycles (e.g., Adams, 1996b; Mörner et al., 2003; Mörner, 2005). On the other hand, a common misconception about paleoseismology relates this discipline to the study of prehistoric earthquakes only, as originally defined by Wallace (1981). This surely used to be, and still is, its main focus, but it does not cover all current potential spectra and unnecessarily restrains paleoseismology to a strict time window. In fact, paleoseismology should be understood as the study of the ground effects from past earthquakes preserved in the geologic and geomorphic record (Michetti et al., 2005b), regardless of the time of occurrence. In a more general way, it aims at studying any geologic deformation related to earthquakes (Audemard, 2005). In that sense, the separation among disciplines studying past individual earthquakes, i.e., instrumental seismology, historical seismology, archaeoseismology, and paleoseismology, should not be based on strict time windows of observation, but on the type of source information that is being used and managed by the researchers. In other words, the distinction between these four different disciplines is fundamentally methodological, both in terms of applied tools and research objects. Consequently, we share the proposal of Caputo and Helly (2008), in which each discipline focuses on different objects: (1) analogue or digital instrumental records, in modern

Geological criteria for evaluating seismicity revisited seismology; (2) oral or written witnesses, in historical seismology; (3) artifacts, in archaeoseismology, where artifact is defined as any product of human activity; and (4) natural features, in paleoseismology. Particularly, “natural object” in this paper is a generic name given to all geologic and geomorphologic features, indicators, evidence, effects, and markers pinpointing the occurrence of an earthquake. The methods and techniques employed for paleoseismic characterization of source parameters are diverse and numerous (for a thorough review, refer to reference textbooks such as McCalpin, 1996; Yeats et al., 1997; McCalpin, 2009) but are still evolving. Methodologies are continuously changing and adapting to encompass very broad and multidisciplinary approaches for the characterization of past earthquakes. Fault trenching investigations have become a cornerstone for paleoseismic analysis because they have the potential to provide a direct assessment of the amount and timing of recent coseismic and total (coseismic + aseismic) fault slip. This has led to a common misconception that trenching and paleoseismology are almost synonymous or equivalent (e.g., Caputo and Helly, 2008; Carrillo et al., 2009). As a matter of fact, trenching is only one technique among many other paleoseismic approaches, with its own specific advantages and limitations. For further details on this issue, we refer to McCalpin (1996), Yeats et al. (1997), Audemard (2005), Michetti et al. (2005b), and McCalpin (2009), among other reviews. The study objects that are the focus of paleoseismic investigations have also evolved with time and have particularly increased in number as a response to new demands in the field of seismic hazard assessment (SHA). From the understanding that a tectonic deformation may induce a sedimentological response, which can then be fossilized and preserved in the geologic record, paleoseismologists, and particularly geologists and tectonic geomorphologists, have continuously developed new paleoseismic objects to derive the timing and magnitude of hazardous earthquakes. These developments in paleoseismology have led to the incorporation of an ever-growing number of earth sciences disciplines, such as seismology, Quaternary geology, tectonics, structural geology, sedimentology, (sequence) stratigraphy, pedology, geomorphology, geochronology, remote sensing and geophysical prospecting methods, such as seismic reflection, side-scan sonar, ground-penetrating radar (GPR), and electric tomography, as well as geodetic surveying and monitoring techniques, such as global positioning system (GPS) and electron distance meter (EDM), among several others. In fact, the present contribution aims at answering the following question that any researcher may pose himself or herself anytime they are requested to assess the seismic hazard of a given fault or region: What object(s) can I study in this region in order to know the level of seismic hazard to which the planned construction project will be subjected? MAIN POTENTIAL STUDY OBJECTS The study and understanding of contemporary earthquakes and their associated effects have provided the basis for the search,

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recognition, and characterization of the types of evidence that paleoseismology needs to unravel from the geologic record. Owing to the physics of earthquakes, two major types of deformations typically occur at ground or sea-bottom surfaces. One set of features relates directly to the earthquake rupture process itself, while other features are the result of the accompanying seismic shaking. Serva (1994), McCalpin (1996), and Michetti et al. (2007) referred to these as primary and secondary evidence, respectively. Depending on their location (or proximity to the causative fault or surface break), McCalpin (1996) subdivided these mechanisms into on-fault (near-field) and off-fault (farfield) evidence. Since the secondary features can occur throughout the entire affected region, including on or across the surface break, Audemard (2005) preferred to subdivide the coseismically induced permanent ground deformations into two groups based on their genesis and degree of association with the earthquake, namely direct and indirect ground deformations. The first set of features is directly related to the fault plane and its kinematics. The second group of features includes both the seismically induced effects (mass instability and soft-sediment deformation, including liquefaction) and all other ground modifications at the local or regional scale, such as uplift and subsidence, warping, buckling, bulging, etc. In order to incorporate the many newly employed paleoseismic study objects or targets, this latter classification needs to be given a wider meaning that does not restrict it to ground deformation only and that provides room to include other indirect types of evidence of earthquakes, such as seismites, which do not imply any ground-surface deformation, but refer to sediment remobilization and redeposition, broken or affected speleothems, truncated bioherms, tree growth changes recorded in dendrochronology samples, tsunamites, and even some types of soft-sediment deformations that include sills and convoluted bedding. McCalpin (1996) added a third level of subdivision based on the timing of the feature with respect to the time of occurrence of the earthquake, namely instantaneous (coseismic) and delayed response (postseismic). The need for this third level of classification does not appear to have any practical application because paleoseismic studies typically can make no distinction between coseismic and postseismic processes. Furthermore, the time taken by delayed processes (e.g., deposition of colluvial wedges, filling of open cracks, remobilization and redeposition of sediments, tree growth recovery or healing, among others) is negligible compared to the recurrence interval between large earthquakes in most faults. We herein propose to subdivide the indirect or off-fault objects into three subgroups or types: Type A encompasses seismically induced effects, including soft-sediment deformation (soil liquefaction, mud diapirism), mass movements (including slumps), broken or disturbed speleothems, fallen precarious rocks, shattered basement rocks, and marks from degassing (mud volcanoes, seeps, pockmarks). Type B consists of remobilization and redeposition of sediments (turbidites, homogenites, and tsunamites) and rocks (e.g., large boulders launched by tsunami waves on rocky coasts, i.e., erratic blocks). Type C includes regional markers of

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Audemard M. and Michetti

tectonic uplift or subsidence, such as flexural-slip thrust faults, reef tracts, (truncated) bioherms, micro-atolls, terrace risers, river channels, progressive unconformities, and mountain fronts. ON-FAULT PALEOSEISMIC OBJECTS On-fault features, i.e., those related to fossil or fresh surface rupture of earthquakes, attracted the attention of researchers from the early days of paleoseismology. Owing to the scarcity of natural exposures of these features (with very rare exceptions, such as the case studied by Audemard et al. [1999] and also the spectacularly preserved surface ruptures of normal fault earthquakes in Nevada [e.g., Slemmons, 1957; Wallace, 1984]), or the rare availability of anthropogenic cuts across them (e.g., Nicol and Nathan, 2001; Mörner et al., 2003; Kuhn, 2005; López, 2006; López C. and Audemard M., this volume), paleoseismic trenching has become the most widespread technique of viewing these ground or sedimentary disturbances in the late Quaternary geologic record. We could actually state that paleoseismology at its early stages essentially developed around trenching studies. Trenching is typically practiced across active fault traces that are currently well identified by means of their geomorphic expression via fault-related landforms (D in Fig. 1A and I in Fig. 1B). Several hundreds of these studies are reported in the literature (e.g., Sieh, 1978; Machette et al., 1987; Crone and Luza, 1990; Van Dissen et al., 1992; McCalpin et al., 1993; Okumura et al., 1994; Audemard, 1996; Michetti et al., 1996; Townsend, 1998; among many others), but even larger numbers of paleoseismic studies are of a confidential nature or can only be found in unpublished reports, particularly in the United States, as many of these studies have been performed in compliance with regulations or legislation. This volume includes several new cases of these types of studies (Audemard, Clark, Lalinde P. et al., López C. and Audemard M., McCalpin, Olig et al.). Paleoseismic trenching strongly relies on a previous very thorough and detailed neotectonic assessment of the fault or region (Audemard and Singer, 1996, 1997, 1999; Audemard, 2005; Michetti et al., 2005b) to identify the best potential trenching sites. The main aim of this technique is to expose the interplay of fault activity with recent sedimentation in trench walls. Young (late Pleistocene to Holocene) sediments record perturbations that are directly related to the fault interference with the sequence. We emphasize that successful results from trench assessments heavily rely on the recognition of active or capable fault traces and the understanding of the interaction between tectonics and sedimentation at the chosen trench site, prior to excavation. The only way of bypassing this lengthy procedure of trench-site selection is when a trench can be directly excavated across a newly ruptured fault surface break (i.e., the Cariaco 1997 earthquake surface rupture studied by Audemard [1999a, 1999b, 2007, this volume], and the ChiChi 1999 rupture studied by Chen et al. [2004]), in which case an accurate location of the seismogenic fault trace(s) and the tectonic style of the active fault are provided. Finally, the trenching approach is not only applied to on-fault features but also to all

those perturbations indirectly induced by active faulting or by seismic shaking that may be suitably recorded in the sedimentary sequence (Fig. 1). This issue shall be dealt with in more detail later in this paper. In particular settings, geophysical investigations have been of great help in trench studies. The most commonly applied geophysical methods are ground-penetrating radar, which is widely known by its acronym GPR (e.g., Cai et al., 1996; Jol et al., 2000; Reicherter et al., 2003; Ollarves et al., 2004; Maurya et al., 2005), and electrical resistivity tomography (ERT; e.g., Silva et al., 2001; Caputo et al., 2003; Colella et al., 2004; Fazzito et al., 2009). These methods have helped to locate the active fault trace with a much higher precision; these traces often exhibit subdued or very subtle earthquake-related geomorphic features or are situated in land that has been heavily modified by anthropogenic activities (e.g., Cushing et al., 2000; Dost and Evers, 2000; Jongmans et al., 2000; Lehmann et al., 2000; Meghraoui et al., 2000; Verbeeck et al., 2000; Demanet et al., 2001). In the particular case of the surface trace of a gently dipping thrust fault that tends to parallel and mimic the topographic contours, the localization of the fault trace solely on the basis of morphologic expression is very difficult and requires the support of these geophysical surveys prior to trenching (e.g., Chow et al., 2001; Fazzito et al., 2009). In addition, GPR and/or ERT have also been applied to onshore strikeslip faults, where often the fault splays into several branches, and these techniques serve to determine the branches that are the most active (e.g., Gross et al., 2000; Wise et al., 2003; Audemard et al., 2006; Kürçer et al., 2008). The most common on-fault geomorphic features used as paleoseismic indicators, from a geomorphologic viewpoint, are fault scarps or counterscarps, sag or fault ponds, pop-ups or pressure ridges (at smaller scale, these are known as mole tracks), shutter ridges, and open fissures, among others. Regardless of the tectonic style (thrust, normal, or strike-slip faulting), fault scarps are definitely the most widely assessed feature due to their common occurrence (D and G in Fig. 1A and I in Fig. 1B). Furthermore, they constitute a perturbation at the ground surface, which triggers morphodynamic processes such as erosion, redeposition, and surface smoothing that have a sedimentary signature and are prone to fossilization. Depending on the balance between tectonic and sedimentation/erosion rates at the scarp, three different situations are recognized that are recorded distinctively: (1) When tectonic rate is higher than sedimentation rate, the scarp grows through time and sedimentation is bounded to the “lowtopography” block; (2) when tectonic and sedimentation rates are similar, available space is filled by sediments and the fault-related landform is leveled; and (3) when sedimentation rate is higher than the tectonic rate, the landform is buried by sediments. This last condition leads to blind normal or thrust faulting, which in turn induces growing fault-propagation folding. Depending on whether the difference between sedimentation and tectonic rates is small or large, this deformation ought to be assessed either as on-fault evidence or an off-fault feature, respectively. In regions where event dating is difficult to impossible due to lack or scarcity

Geological criteria for evaluating seismicity revisited of datable materials (K in Fig. 1B), techniques based on the diffusion equation (e.g., Begin, 1992; Hanks, 2000) have been applied to date scarps by estimating the degree of scarp degradation. This technique experienced a boom in the western United States in the 1970s and 1980s (e.g., Bucknam and Anderson, 1979; Nash, 1980; Colman and Watson, 1983), but its generalized application on a worldwide scale has become very limited because it is strongly dependent on local climate, which may vary quite rapidly within a basin or along a mountain front. It definitely depends on a local diffusivity (Phillips et al., 2003), which is different from place to place. In recent years, cosmogenic dating (radionuclides 10 Be, 26Al, and 36Cl, and stable nuclides 3He and 21Ne) of fault scarps by exposure to cosmogenic radiation has largely replaced the scarp degradation technique. This approach has been applied to fault scarps both in Quaternary alluvial deposits (e.g., Phillips et al., 2003; Siame et al., 2006) and bedrock (e.g., Zreda and Noller, 1998, 1999; Mitchell et al., 2001; Siame et al., 2006). Most commonly, this technique provides data on tectonic slip rates during the time span of scarp exposure to cosmogenic radiation (e.g., Ritz et al., 1995; Van der Woerd et al., 1998; Palumbo et al., 2004). In a similar manner, it has been used for dating nontectonic scarps, such as sackung uphill-facing scarps (B in Fig. 1A) by Hippolyte et al. (2006), which could also eventually provide pertinent paleoseismic information in the cases where these large mass instabilities were seismically triggered (refer to section “Seismically Induced Effects” later herein). As mentioned earlier, scarps are the features most commonly trenched for paleoseismic purposes. The success of this type of investigation is enhanced if sag- or fault-bounded ponds form against them. However, tectonic scarps are seldom preserved in the case of thrust faults, unless the fault cuts across very competent rocks. Most Quaternary thrust faults cut across soft or unconsolidated deposits, since they tend to crop out at the foot of the mountain front. It is frequently the case that tectonic slip rates of such thrust faults are much less than the sedimentation rate, which can lead to an almost complete masking of the active thrust fault by younger sediments, creating the situation of a blind thrust fault. If no other option is available, paleoseismologists have overcome this situation by trenching sympathetic faults—secondary faults that form in mechanical and kinematical association with the main thrust fault, such as scarps of flexuralslip thrust faults (Fig. 1B; e.g., Yeats, 1986a, 1986b; Costa et al., 1999; Livio et al., 2009). Back-thrust faults are also a potential trenching target. They happen to exhibit a higher dip than the main thrust and very frequently face in an upstream direction, which offers a better scenario for the interaction between tectonics and sedimentation. The major limitation in these situations is the low degree of certainty the researcher has on the actual kinematic activation of the sympathetic and/or antithetic faults in relation to the coseismic slip of the seismogenic thrust fault. In other words, the sympathetic fault does not necessarily move each time the main thrust fault does, which converts it into a poorly reliable chronometer. Moreover, although mechanically connected, there are as yet no published cases of the relation-

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ship between actual slip on the major fault and the induced slip on the flexural-slip faults. Earthquake size estimates in this case should be based on: (1) intensity data (Serva, 1994), which are then converted into magnitude values, when available, and (2) the application of the concept of seismic landscape (Michetti et al., 2005b). The definition of the relationship between macroseismic intensity and earthquake ground effects was in fact the rationale for the introduction of the ESI 2007 Intensity Scale proposed by INQUA (Michetti et al., 2007; Reicherter et al., 2009). Generally speaking, all the previous geomorphic objects, which favor the continuous fine-grained, thinly bedded sedimentation against the fault plane, constituting the ideal setting for the recording of future ground-surface deformations, have proved to be really useful for paleoseismic studies. After overcoming water problems (through draining/water pumping in tropical regions and digging frozen ground in temperate regions), sag and fault ponds have generally provided the best and most complete paleoseismic records (e.g., Audemard, 1997; Lindvall et al., 2002; Audemard, this volume). Several favorable geologic conditions (for more details, refer to Audemard, 2005) have to come together at these “trenchable” sites in order to increase the chances of successful seismic hazard assessment. The paleoseismic assessment of any of the aforementioned objects can yield information on most or all of the following aspects (Audemard, 2005): confirmation of Holocene fault activity, slip per event and average slip rate of a given fault or fault segment, slip vector decomposition from fault-plane kinematic indicators, recurrence intervals and magnitude of the larger earthquakes (seismic potential) on known faults, coseismic rupture length, fault segmentation, fault interaction as a consequence of stress loading by stick-slip on contiguous faults, time-space distribution of seismic activity along a given tectonic feature, elapsed time since latest event and estimated time to the next event, which determines the likelihood of occurrence of a future earthquake, seismotectonic association of historical earthquakes, and short- and long-term landscape evolution. In cut or trench exposures, the most frequent kinds of geologic evidence of the interaction of sedimentation and faulting (Amit et al., 1995; Audemard, 2005), which are interpreted as the result of individual earthquakes, are: (1) filling of groundsurface open fissures (tension or T cracks, synthetic Riedel or R shears, or their combination); (2) fault-scarp–bounded deposition on a downthrown block; (3) a colluvial wedge derived from scarp degradation/subduing, occasionally interspersed with sedimentation on the downthrown fault compartment; (4) event horizon (faulting sealed by a leveling depositional episode); and (5) normal-fault propagation folding, later buried by overlying beds. Trees have been excellent time recorders—chronometers— of the occurrence of large contemporary earthquakes, in association with, as well as away from, their surface breaks (e.g., Sheppard and Jacoby, 1989). The study of the annual growth rings of certain particular plant species, known as dendrochronology, which is somewhat equivalent to varve chronology in lakes of temperate regions (which has applications in higherlatitude regions of the globe that exhibit a marked seasonality),

Figure 1. Generalized sketches of different potential settings of coexisting natural indicators of seismic activity (see text for explanation). (A–G) Upper cartoon depicts a subduction setting, where the paleoseismic geomarkers in relation to the marine environments are to the left, while those related to the orogen are to the right. The lower cartoon emphasizes the objects related to several subaqueous subenvironments. (H–L) The upper diagram tries to show the potential paleoseismic objects existing in a compressional environment, dominated by either outcropping or blind thrust faults. The lower diagram summarizes the paleoseismic objects to seek in a cratonic environment subject to glaciation-deglaciation processes, such as Scandinavia. LGM—Last Glacial Maximum.

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Figure 1. (Continued.)

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has revealed that certain growth changes (ring eccentricity, dead and healing zones, among others) that are the result of seismically induced perturbations can be recorded in big, long-lived tree trunks (e.g., Meisling and Sieh, 1980; Yamaguchi et al., 1997). In response to the natural positive phototropism exhibited by most large trees, loss of tree verticality—tilting of a tree—is soon naturally repaired in the years directly after the perturbing event, which is nicely recorded in the growth of annual rings. This has been shown in localities close to or right on top of contemporary earthquake ruptures (e.g., Page, 1970; Sheppard and White, 1995; Jacoby et al., 1997; Lin and Lin, 1998; Carver et al., 2004). In some cases, disruption of tree roots by ground cracks (e.g., Tasdemiroifelgülu, 1971; Carver et al., 2004) is marked by a deceleration of ring growth, which records the partial or total death of the tree (e.g., Toppozada et al., 2002; Doser, 2004). Even more astonishingly, large trunks, especially of sequoia trees, have been split into two by a crosscutting earthquake surface rupture (G in Fig. 1A; e.g., Carver et al., 2004). This illustrates that growth anomalies of tree rings can be considered as direct or indirect evidence of surface faulting, depending on the degree of association with the ground rupture. The limitation of this kind of evidence, as seen from a paleoseismic viewpoint, resides in the low probability of its preservation and the even lower probability of actually finding the preserved evidence. It has certainly proved to be useful as long as the tree lives. In that sense, this technique in very recent times is being used widely for mass movement monitoring and detection (e.g., Pop et al., 2008; Surdeanu et al., 2008), although its applicability has been known for quite a while (e.g., Terasmae, 1975; Hupp et al., 1987). In coastal areas along subduction zones, it has been observed that after the occurrence of large subduction earthquakes, large coastal areas have drowned after the seismic energy was released, as evidence of coseismic subsidence, due to elastic rebound. This has led to the death of extensive stands of trees as a result of the flooding by marine-water incursions of the low-lying areas and is commonly referred to as “drowned or ghost forests” (e.g., Atwater, 1987; Hamilton et al., 2005). This type of event is preserved in the sedimentary record as a couplet consisting of a horizon of organic-rich terrestrial sediments overlain by a marine sediment layer (A in Fig. 1A; e.g., Nelson et al., 1996; Cundy et al., 2000). OFF-FAULT PALEOSEISMIC OBJECTS The term “off-fault evidence” as defined by McCalpin (1996), or “indirect perturbations” as originally defined by Audemard (2005) and enlarged herein, covers all the earthquake-related or earthquake-triggered evidence different from ground-surface rupturing coinciding with the surface trace of the causative fault, although they can occasionally concur with the fault plane as nearfield features. To illustrate this eventual coalescence of processes, a sand dike, which results from earthquake shaking that triggers liquefaction, takes advantage of the recently displaced causative fault plane, as has been observed in trench walls by Rockwell et al. (2001) and Audemard et al. (2008), because it probably has

a transient higher permeability that facilitates the venting process several seconds to a few minutes after the earthquake occurrence. Since the off-fault evidence is commonly not associated with the causative fault plane, its paleoseismic interpretation needs more elaboration than the on-fault objects. For instance, the interpretation of these features in terms of earthquake magnitude is rarely unequivocal, unless their evaluation is carried out spatially (e.g., Ricci Lucchi, 1995). Each object of the same nature or type (e.g., slide, sand blow, or broken speleothem) should be envisaged as a single “seismometer,” and only the integration of them in a “network” relation can provide a reliable estimate of earthquake magnitude. This tacitly implies that all objects of the same type need to be ascribed to the same event. They must be temporally related; otherwise they cannot belong to a network. Significant progress has been achieved by this approach for mass movements and seismically induced liquefaction. Using modern earthquakes as analogues, relationships between earthquake magnitude and size or spatial distribution of slides have been derived (e.g., Youd and Hoose, 1978; Keefer, 1984; Wilson and Keefer, 1985; Keefer and Manson, 1989; Rodríguez et al., 1999; Crozier, 1992; Keefer, 2000) that can be applied, as an inverse method, to fossil landslides in paleoseismic studies to estimate the magnitude and localization of the epicenter of the triggering event, and by association, the causative fault as well. The same procedure has been carried out for liquefaction features (e.g., Kuribayashi and Tatsuoka, 1975; Youd and Perkins, 1987; Ambraseys, 1988; Barlett and Youd, 1992; Papadopoulos and Lefkopoulos, 1993; Munson et al., 1995; Allen, 1986; Galli, 2000; Castilla and Audemard, 2002; Rodríguez et al., 2002; Rodríguez Pascua et al., 2003; Papathanassiou et al., 2005; Rodríguez et al., 2006; Castilla and Audemard, 2007). On the other hand, and this is a very sensitive issue, the size of these objects may have no relation with the earthquake magnitude or the seismic attenuation laws. For instance, the lateral extent and thickness of homogenites, and tsunamites, tend to reflect pre-earthquake local conditions of the area where they occur. This is also the case for sand dikes formed in association with lateral spreading (Fig. 1A). Their final width, contrary to claims made by Obermeier (1996), is strongly conditioned by the size of the nearby available void (both in width and height) that sets in motion the horizontally sliding mass, creating the fractures through which venting will occur. In many cases, these offfault features provide a good time constraint for the occurrence of the inducing earthquake. So, they can be considered as reliable chronometers. The features implying redeposition appear to be best suited in this respect. These include tsunamites, homogenites, and erratic (tsunami-transported) boulders and blocks, but also broken speleothems, affected bioherms, mass movements, particularly those with ground-surface disruption, and even sand blows. In fact, their occurrence gives a good indication of the timing of the earthquake. On the other hand, other objects can merely reveal that an area is seismo-tectonically active, only due to their occurrence within the sedimentary record. Many soft-sediment deformations, even after reliably confirming their seismic origin,

Geological criteria for evaluating seismicity revisited fall into this category. All those soft-sediment deformations that maintain parallelism to the stratification after remobilization, including convolute bedding, contorted layers, basal surface of slumped beds, flame structures, ball and pillars, sand-vented sills, etc., are useless in terms of earthquake chronometers. Also, penetrative liquefaction features, such as sand dikes, like the ones depicted by Rodríguez-Pascua et al. (2000), need careful handling in terms of age determination of the triggering earthquake. If there is no convincing proof that such a dike reached the preearthquake ground surface by being connected to a sand blow at the top, the age of the youngest disrupted bed should be considered as the oldest age of occurrence of the associated earthquake. In other words, the age of the event in such a case is younger than the last affected stratigraphic marker. This is also valid in the cases when the sand-venting dike is later truncated by erosion. Based on their generating mechanism, in combination with the paleoseismic approach that could be applied to their study, we have subdivided the indirect or off-fault objects into three subcategories: Type A includes seismically induced effects; type B consists of remobilization and redeposition of sediments and/or rocks; and type C entails regional markers of vertical deformation (surface level changes). These are the features that have been the subject of significant development in the past years; in some cases, huge progress has been made. Paradoxically, the inaccessibility of subduction fault zones can be primarily held responsible for the recent progress in the study of all these indirect types of earthquake evidence, as anticipated by Adams (1996b). The reality that subduction zones lie underwater and are responsible for the largest known and most destructive contemporary earthquakes worldwide (e.g., 22 May 1960, Valdivia; 27 March 1964, Anchorage; 26 December 2004, Sumatra; and 27 February 2010, Concepción, earthquakes), in combination with their very large extent around the Pacific rim, as well as their position in relation to large population centers, with their costly infrastructure along the Pacific shores, make the assessment of their seismic potential an urgent necessity. This all has led to the realization that it is necessary to study ground deformations and on-land sedimentary disturbances as evidence of earthquakes in the coastal zones (e.g., Atwater et al., 1995), as well as along the submarine trench during earthquakes. Adams (1996b) had already conducted a rather thorough overview of the ways in which the seismic hazard assessment of such a complex active tectonic and geodynamic setting could be tackled, by presenting the case of the West Canadian coast. A multiproxy investigation, using several of the potential indirect objects (Fig. 1), actually appears to be the most appropriate approach in order to derive the seismic history (earthquake chronology) of a given subduction segment or a given region. It includes any or some of the phenomena, such as earthquake-triggered liquefaction, coseismic land-level changes, and associated sedimentary response (interbedding of marine and continental deposits, growth increase or truncation of living [brain] corals, emerged or submerged coseismic marine terraces), and onshore tsunamites, as well as underwater mass wasting, pockmarks from degassing/dewatering, mud and sand

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volcanoes, offshore turbidites, and homogenites. This concept of a multiproxy approach has been applied to continental regions as well, for instance, in the interior of Canada (Adams, 1996b), in the alpine lakes of Switzerland and Italy (Becker et al., 2005; Fanetti et al., 2008), and in Venezuela (Audemard, 2005). Eventually, by applying some of the existing relationships treated in the literature for mass wasting and/or liquefaction, mentioned earlier, we will be able to make a minimum estimate of the triggering paleo-earthquake size. Seismically Induced Effects Mass movements and ground liquefaction are the two most common and widespread natural phenomena associated with earthquakes worldwide. These effects, induced by earthquake shaking, have been commonly recognized and described for onshore environments in historical documents over the centuries. For instance, description of liquefaction features by Sarconi (1784) and illustration of mass movements by Simón (1627) are classical examples of historical reports for seismically induced geological effects. Mass movements, because of their requirement of an energy gradient, typically occur in mountainous regions, but they can also affect large flat areas in the form of lateral spreading (Fig. 1A). Conversely, liquefaction tends to affect geologically young, low-lying, flat alluvial fill areas of fluvial or coastal environments. However, these two phenomena are not exclusive of onshore environments. On the contrary, they appear to develop underwater, in sea-bottom and lake-bottom sediments, and are more frequent and more widespread, as observed during contemporary earthquakes (e.g., Sims, 1973, 1975; Mosher et al., 1994; Nakajima and Kanai, 2000), as well as known from the past geologic record (e.g., Hempton and Dewey, 1983; El-Isa and Mustafa, 1986; Ringrose, 1989; Roep and Everts, 1992; Maltman, 1994; Hampton et al., 1996; Marco et al., 1996; Mörner, 1996; Rodríguez-Pascua et al., 2000; Mörner et al., 2000, 2003). These induced effects have proved, during many destructive earthquakes, to be more harmful than the earthquake shaking itself. As an illustration of this, there is the large destruction produced by the Prince William Sound, Anchorage-Alaska, 1964, earthquake at Turnagain Heights due to lateral spreading (Hansen, 1965). In the same way, either large mass-wasting bodies (e.g., Ancash, Perú, 1970, earthquake, with 18,000 casualties; Cluff, 1971; Plafker et al., 1971) or widespread mass movements have caused generalized destruction and numerous fatalities in urban areas, such as during the Deixi, Sichuan-China, 1933, earthquake, with 6800 casualties (Li et al., 1986). Associated with slope instabilities, an even more harmful factor has been the breaching of slide-dammed lakes days or months after the earthquake (Fig. 1B; e.g., the Mocotíes 1610 event in the Mérida Andes—Singer, 1998; Ferrer, 1999; the Deixi, Sichuan-China, 1933, earthquake, with 2500 casualties—Li et al., 1986). As for outcropping active fault planes, trenching has been practiced onshore for both earthquake-induced gravitational scarps (B in Fig. 1A; e.g., Wallace, 1984; Crozier, 1992; Nolan

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and Weber, 1992, 1998; McCalpin, 1999; Onida et al., 2001; McCalpin and Hart, 2002; Tibaldi et al., 2004; Gutiérrez et al., 2005, 2008) and liquefaction features (e.g., Tuttle et al., 1990; Audemard and De Santis, 1991; Tuttle and Seeber, 1991; Clague et al., 1992; Walsh et al., 1995; Tuttle, 2001; Mörner et al., 2003; Cox et al., 2004; Guccione, 2005) in many different tectonic settings. Not only have these features proved to be the only means of assessing the seismic history of subduction zones (A in Fig. 1A), but they have been used elsewhere where faults have not ruptured the ground surface. For instance, trenching for liquefaction features has been carried out in the New Madrid area to recognize not only the effects of the 1811–1812 earthquake sequence and their extent, but also those of its precursors (Tuttle, 2001; Cox et al., 2004; Guccione, 2005). In cratonic areas subject to glaciation and deglaciation processes during the Quaternary, such as northern North America and Scandinavia, where Holocene sediments act as young recorders of recent deformation, and recognized onshore (intraplate) faults with Holocene surface rupture are scarce (e.g., Lagerbäck, 1979; Adams, 1989; Grant, 1990; Adams et al., 1991; Fenton, 1994), the study of indirect paleoseismic objects, as well as other features or approaches (e.g., tsunami deposits, age determination of fault-scarp exposure and scarp-profile degradation studies, dendrochronology, lichenometry; Fig. 1B), has clearly established that these socalled stable regions are seismically active. Adams (1996b) and Mörner (2005) have found that unexpected contrasting seismic patterns at the time of deglaciation, characterized by large and very frequent earthquakes, also took place. In addition, the study of earthquake-induced effects (mass wasting and soft-sediment deformation in a more general way) is becoming a complementary tool to on-fault trench studies. In Venezuela, several periglacial (paleo-)lakes directly offset by the active Boconó fault have been cored and sampled in order to demonstrate the applicability of this approach, with the intention of extending its use to seismically active regions with no outcropping seismogenic faults (Carrillo, 2006; Carrillo et al., 2006a, 2006b, 2009). In Israel, very long earthquake histories have been reconstructed through the study of soft-sediment deformations in late Pleistocene to Holocene laminated lacustrine sequences of the Dead Sea region (Marco et al., 1996; Agnon et al., 2006). In addition to directly exposing the seismic evidence through trenching, these past seismically induced effects in the geologic record have been assessed through indirect methods, either invasive (coring and geotechnical methods; E in Fig. 1A and H in Fig. 1B) or noninvasive (geophysical methods). Fossil liquefaction features (particularly sand blows) in the subsurface have been uncovered by GPR (e.g., Liu and Li, 2001). Since these earthquake-induced effects may also occur underwater, their investigation under such conditions has mostly relied on the combination of geophysical methods and sediment core recovery (Figs. 1A and 1B). The overall geometry of the sedimentary bodies and their stratigraphy are obtained from the geophysical investigation, which can call upon methods such as side-scan sonar to provide highly detailed sea-bottom topography, high-resolution shallow seismic, including multi-

beam, sparker, boomer, and pinger (subbottom profiler), among others, while coring reveals the physical characteristics of the sediments and deformations, and also brackets the time of occurrence of such effects (if dated). These underwater objects have been studied both at sea (Schafer and Smith, 1987; Piper et al., 1999) and in lakes (Beck et al., 1992; Shilts and Clague, 1992; Thomas et al., 1993; Beck et al., 1996; Chapron et al., 1996; Becker et al., 2005; Schnellmann et al., 2005; Carrillo, 2006; Carrillo et al., 2006a, 2006b, 2009; Beck, this volume), employing some of the aforementioned methods. The timing of activation (or reactivation) of earthquakerelated mass movements has been derived from trenching across the scarp or scar at the slide crown of such mass wasting (e.g., Wallace, 1984; Onida et al., 2001; McCalpin and Hart, 2002; Tibaldi et al., 2004; Gutiérrez et al., 2010). In this sense, large sackungen have received particular attention (B in Fig. 1A; e.g., Gutiérrez et al., 2005, 2008). From these studies, it is noteworthy that a significant uncertainty remains as to the completeness of the earthquake chronology derived by these methods. It still remains a difficult task to ascertain that a particular slide of a given size and certain characteristics will always be prone to slide during an earthquake of a particular magnitude or intensity above a certain threshold. It is equally difficult to assign a seismic origin to many sackungen in high mountain regions, especially if their occurrence is concurrent with deglaciation periods (Hippolyte et al., 2006). If the displacement of the mass movement is large enough and the mobilized mass does not completely fall down, the consequent depression generated at the foot of the scarp may create new space for the emplacement of small lakes. In such a case, the earthquake chronology can be assessed by the study of the perturbations recorded in the bottom sediments of these lakes (B in Fig. 1A; soft-sediment deformation, turbidites, [seicheinduced] homogenites, slumps, among others). If these lakes or ponds happen to be small in size, only coring can be performed. If a lake is large and deep enough, geophysical prospecting can be undertaken. Some large mass movements, particularly rock avalanches, but also huge debris and mud flows, may have important runoff and may block the valleys by their downstream deposits. This may result in the damming of water bodies of considerable volume upstream. This process has been observed in several historical and contemporary earthquakes (e.g., Simón, 1627; Li et al., 1986; Clague and Evans, 2000). Consequently, the recognition of the aforementioned markers (soft-sediment deformation, turbidites, seiche-induced homogenites, slumps, among others) in lake-bottom deposits in slide-dammed lakes has also been used as a natural indicator of past earthquakes (Costa and Schuster, 1988; Weidinger, 1998; Hermanns et al., this volume). This same approach can be applied to water bodies that are pounded against a coseismically growing active fold (H in Fig. 1B). An example of this happened at Oued Foda during the 1980 El Asnam–Algeria earthquake, above an active blind thrust fault (Philip and Meghraoui, 1983). The analysis of such lake fills may help to reconstruct the seismic history of these faults

Geological criteria for evaluating seismicity revisited that show no evidence of brittle expression at the surface as direct evidence. Geophysical methods, such as vertical electrical soundings (VES) and ERT, have been used jointly to calibrate multielectrode profiles and to estimate the depths of landslide deposits in the town of Celano (Fucino Basin, central Italy; Rinaldini et al., 2008). Assuming that the quantum leaps forward in multichannel seismic acquisition technology are maintained in the years to come, we could easily envisage a situation in which much higherresolution three-dimensional (3-D) seismics (seismic cube) would allow the regional-scale mapping of most of these seismically induced effects, particularly those paleoseismic objects that rest on a perturbed ground surface or seabed (e.g., sand blows, slides, but also mud volcanoes and pockmarks; Fig. 1). This modern seismic method currently allows the construction of time slices, although not yet of sufficiently high resolution. These constitute a reconstruction of the environment or topography at a given time and thus would fulfill the same objectives as current macroseismic surveys, from which magnitude and epicenter of the causative earthquake could be derived by applying known relationships. In other words, at each identified event horizon, corresponding to a time slice in the seismic cube, the paleoseismic objects would be mapped in as much detail as the resolution of the method would permit. Those phenomena described here are known in the literature as seismically induced effects, but other, less well-known, natural features also form as a result of seismic shaking, among which we shall discuss the following: mud volcanoes, oil seeps and pockmarks, broken speleothems, and toppled precariously balanced rocks. Mud volcanoes/diapirs (F in Fig. 1A), and even salt domes, can be used as analogs for earthquake-triggered sand venting in the form of dikes connected to sand blows. Although it is known that mud diapirs may form and flow to the surface during earthquakes (e.g., Arnold et al., 1960), their use as paleoseismic indicators, to our knowledge, has not yet been reported in the literature. This can be attributed to the fact that they are not exclusively associated with earthquakes. Furthermore, many of these features do not have an episodic motion, but rather they have a steady-state ascending motion, which would make it more difficult to ascribe a seismic origin, although their upward motion and also final eruption are known to accelerate during earthquake shaking. The process of seabed degassing or gas seepage, which produces pockmarks as surface evidence (Hovland and Judd, 1988; Hovland et al., 2002; Pinet et al., 2008) is also known to accelerate during earthquakes (Fig. 1A). Pockmarks are shallow depressions (5–300 m in diameter and 2–20 m deep) formed by the explosive release of gas (Gluyas and Swarbrick, 2004). Seismic shaking can provoke unusually large flows from gas or oil seeps (Levorsen, 1967; Field and Jennings, 1987), similar to what happens with water springs that may resume or stop flowing as a consequence of earthquake action (e.g., González et al., 2004). For instance, following the 1971 San Fernando Valley (California) earthquake, several of the previously inactive oil seeps in the area resumed activity (Mandel, 2001). Unfortunately, there is

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no way that degassing evidence will be recorded in the geologic record, but pockmarks are perfectly well preserved on the seabed, as revealed by side-scan sonar. These features potentially provide the same paleoseismic information as earthquake-triggered sand blows. Underwater, they can be detected by high-resolution single-channel seismics (sparker or subbottom profiler) or multibeam surveys, whereas age bracketing can be obtained on coring. So far, it has not been applied for specific paleoseismic purposes. Speleothems are also excellent recorders, as well as highly precise chronometers of seismic shaking (C in Fig. 1A). For over 20 yr, speleothems have been used as indicators of seismic activity (Postpischl et al., 1991; Bini et al., 1992; Gilli, 1996, 1999; Delaby, 2001; Pérez López et al., 2009). This has been stimulated by the fact that they are easily datable (by 14C, for instance), they record growth perturbations very neatly, and they also grow by annual precipitation, similar to tree rings. Their principal limitation is that their occurrence is restricted to soluble carbonatic rocks. Anyhow, their steady way of growing permits detection of growth pattern disturbance as a result of catastrophic events with a temporal resolution of only 1 yr, similar to varve chronology and dendrochronology. Because of these characteristics, speleothems have been an object of many paleoseismic studies, as confirmed by the large number of studies in which historic (Gilli, 1999; Lemeille et al., 1999) and prehistoric (Postpischl et al., 1991; Delaby, 2001; Lignier and Desmet, 2002; Lacave et al., 2004; Kagan et al., 2005) earthquakes have been dated or constrained. Good overviews on the use of speleothems as paleoseismic indicators have been published by Forti (2001), Gilli and Delange (2001), and Becker et al. (2006). However, their interpretation is not straightforward, as demonstrated by Cadorin et al. (2001) and Gilli (2004). The latter author indicated that the breakage or failure of speleothems may also be caused by other processes different from seismic shaking, such as glacial advances or retreats. The possibility that sliding of the limestone masses containing the caves occurred as a result of gravitational processes also needs to be ruled out before ascribing a seismic origin, unless sliding can be assumed to have been triggered by a given earthquake. Precarious rocks or precariously balanced rocks indicate that strong earthquake motions have not occurred at a site since their particular precarious situation developed. In that particular negative sense, these rocks or groups of rocks can be effectively used as earthquake strong ground-motion seismoscopes that have been operating on solid rock outcrops for thousands of years (Brune and Whitney, 1992; Brune, 1994, 1996, 1999; Shi et al., 1996; Brune et al., 2003). Thus, estimates of the threshold ground acceleration for the toppling of these rocks can provide constraints on the peak ground accelerations experienced during previous earthquakes. Consequently, the presence of precarious rocks is an indicator for the absence of earthquakes above a certain magnitude. Preservation of these rocks in their precarious position is thus a function of the seismicity of the region under study with a propensity for the larger events, and this reduces its applicability to high-seismicity regions of the world. On the other

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hand, toppled precarious rocks can be used as chronometers of the latest earthquake, the magnitude of which can be estimated as having exceeded the threshold value needed for toppling. The age of the place on which the precariously balanced rocks used to stand (commonly made of hard rock) could be determined by cosmogenic dating, thus providing time of occurrence of the latest event. Likewise, the age of the precarious rock toppling could be derived from dating the age of the ground surface on which the block fell (similar to J in Fig. 1B). If it fell on soil, the 14C method of dating could be applied, whereas if the block has fallen on hard rock or on sediments, even luminescence (thermoluminescence or optically stimulated luminescence) techniques could be applied. Remobilized and Redeposited Sediments We propose to include in this type of indirect paleoseismic object all those deposits of any grain size that have been transiently mobilized by the earthquake or associated phenomena (e.g., underwater or subaerial mass wasting, tsunami waves) and redeposited shortly after. Usually, this type of deposit, which includes turbidites, homogenites, and tsunamites, is generated in aqueous environments but results from different processes or intervening mechanisms. Seismically induced turbidites and homogenites require that underwater (eventually, subaerial) sediments were set in motion by the earthquake shaking (e.g., Francis, 1971; Siegenthaler et al., 1987; Shiki et al., 2000). Turbidites are generally the result of submarine mass-wasting processes causing hyperpycnal inflow that generates turbidity currents, which transport sediments downslope along the seabed (e.g., Adams, 1990, 1996a; Doig, 1991; Inouchi et al., 1996; Gorsline et al., 2000) in the days or months following an earthquake, as evidenced in the case of several twentieth-century earthquakes (e.g., Heezen and Ewing, 1952; Thunell et al., 1999). Homogenites, instead, would rather require the oscillatory motion (seiche) of a confined water body, although this motion is also fed from underwater mass wasting (e.g., slumps; e.g., Beck, this volume) or nearshore subaerial mass movements (e.g., collapse of the edifice of the Santorini Volcano as proposed by Cita et al., 1996). Conversely, most tsunamites are fed from nearshore sediments that are carried inland by tsunami waves (e.g., Atwater, 1987; Clague and Bobrowsky, 1994a, 1994b; Clague et al., 1994; Chagué-Goff and Goff, 1999; Bourgeois et al., 1999; Sawai et al., 2008; among many others). This does not exclude the situation where tsunamites can also appear as homogenites in deep and/or confined water bodies (e.g., Kastens and Cita, 1981; Beck et al., 2007). Typically, tsunamites are thought to be fine- to medium-grained deposits (e.g., Dawson et al., 1988; Shi et al., 1995; Paris et al., 2007), but in fact, the size of the transported material is a function of the tsunami wave energy and availability of materials in the nearshore environment in the case of onshore tsunamites (for a review, refer to Scheffers and Kelletat, 2003), or in the sea bottom for deep-water tsunamites (homogenites). Tsunami waves have been capable of removing and transporting large boulders

inland (e.g., Bourrouilh-Le Jan and Talandier, 1985; Jones and Hunter, 1992) and huge (several m3 in size) blocks of coral reefs, eolianites, or other nearshore deposits (Scheffers, 2004; Scheffers and Kelletat, 2005; Scheffers and Scheffers, 2007). Onshore tsunamites can be assessed in outcrop or by shallow coring (A in Fig. 1A). The coast of the Cascadia subduction zone of the northwestern United States has been a natural laboratory for over the past 40 yr, during which their study has much evolved (Atwater et al., 1995). In the early days, recognition of the tsunamites in that area relied mostly on sedimentary features recognized with the naked eye. The identification of these marine incursions nowadays relies on the combination of several disciplines: sedimentology, geochemistry, paleontology, including micropaleontology, malacology, palynology, and geochronology, among others, which has enhanced and facilitated their recognition in the geologic record. Since all other paleoseismic objects of this sort are mostly preserved in still-active natural environments, particularly in lake- or sea-bottom sequences, they require an integrated paleoseismic assessment relying on joint subaqueous geophysical prospecting methods and piston coring (e.g., Beck et al., 1996; Cita and Rimoldi, 1997; Doig, 1998; Carrillo et al., 2006b; Beck et al., 2007; Beck, this volume). However, paleoseismic indicators of this subset have also been retrieved from outcropping lake sequences (e.g., Doig, 1991; Carrillo et al., 2006a), which unfortunately have been truncated and as a consequence provide only incomplete paleoseismic records for the recent time window. Paleoseismic Indicators of Vertical Motion Most of the indirect paleoseismic objects discussed here constitute an invaluable help when direct evidence of coseismic faulting is unavailable. In other words, their recognition has been greatly enhanced by their study in subduction-related seismogenic source regions and occasionally also in association with blind thrusts. Earlier, we explained how elastic rebound acts coseismically in subduction zones and can be recognized by the presence of the so-called ghost or drowned forests after the occurrence of large subduction earthquakes in the coastal zone, where it is clearly recorded in the onshore sediments by a couplet of organic soils capped by a horizon of marine sediments (A in Fig. 1A). In addition, it can be perfectly recorded by marine biogenic constructions such as coral reefs and (micro-)atolls on the foreshore side of the coastal bulge (Fig. 1B). In inverse relation to the onshore sediment couplet, which registers emergence during the interseismic period followed by subsidence during the earthquake, the coral growth offshore first indicates steady subsidence during the interseismic period followed by abrupt uplift during the seismic energy release (Taylor et al., 1987; Sieh et al., 1999; Zachariasen et al., 1999, 2000). This sudden uplift, when of sufficient magnitude, can result in aerial exposure and subsequent coral head truncation by wave erosion or natural death of the coral species. Similar to tree rings and varves in temperate regions, growth of corals indicates seasonality in tropical regions, and

Geological criteria for evaluating seismicity revisited coral growth is registered in annual bands (e.g., Knutson et al., 1972; Woodroffe and McLean, 1990), but their occurrence is of course limited to the warm regions of the world. Furthermore, certain species of the coral genera Porites and Goniastrea are sensitive natural recorders of lowest tide levels, which make them ideal natural instruments for measuring emergence or subsidence relative to a tidal datum (Scoffin and Stoddart, 1978; Taylor et al., 1987). These Porites and Goniastrea coral heads grow radially upward and outward until they reach an elevation that allows their highest corallites to be exposed to the atmosphere during lowest tides. Consequently, rather small changes in elevation of the coral substratum or sea level induce very clear changes in the growth patterns of these coral colonies or micro-atolls (Zachariasen et al., 2000; Natawidjaja et al., 2007). Thanks to the excellent temporal and vertical resolution of these micro-atolls, it has been possible to decipher the seismic history of particular segments of subduction zones in tropical regions (e.g., Sieh et al., 2008). In the past 30 yr or so, the development of new, modern, and more precise dating techniques such as uranium series, thermoluminescence (TL), optically stimulated luminescence (OSL), and electron spin resonance (ESR), among others, in combination with a more accurate reconstruction of the history of Quaternary sea levels based on variations in the oxygen isotopic composition of seawater (Chappell, 1974; Chappell et al., 1996a), has stimulated and accelerated the study of both emerged and submerged coastal marine features in tectonically active regions worldwide, as well as improved their temporal resolution (Fig. 1B). This has led to an improved determination of vertical slip rates based on marine terraces, either constructional (biogenic or detrital) or erosional. Many appropriate areas have been studied in that respect, and some regions of the world have actually become natural laboratories, such as Huon Peninsula in Papua New Guinea, where the matching of these marine terraces to sea-level curves (Chappell, 1974; Chappell et al., 1996a) has allowed the recognition of other marine terraces, the origin of which has been ascribed to coseismic vertical tectonic motion (Ota et al., 1993; Chappell et al., 1996b; Ota and Chappell, 1996). Other regions of the world have been assessed likewise, such as Chile by Vita-Finzi (1996), New Zealand by Pillans and Huber (1995), and the Gibraltar Strait zone by Zazo et al. (1999). The Ventura Avenue anticline in California deserves particular mention in this respect, where Stein and Yeats (1989) established that this active fold is repeatedly raised by earthquakes recurring roughly every 600 yr, as registered in a series of coseismically uplifted marine terraces. The equivalent of marine terraces in coastal areas, alluvial terraces of the inland environment have been occasionally assessed for the purpose of determining coseismic vertical or horizontal deformations (e.g., Wellman, 1955; Suggate, 1960; Lensen, 1968; Lensen and Vella, 1971; Van Dissen et al., 1992) in continental interiors. However, the most notable contribution of these flights of staircase-style marine, lake, or alluvial terraces is in permitting us to arrive at reliable estimates of the vertical slip rate of tectonic blocks or folds that have been uplifted or subsided.

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CONCLUDING REMARKS Over the past 40 yr, there has been a steady increase in the number of natural objects (generic name herein given to geologic and geomorphologic evidence, indices, markers, effects, indicators, etc.), both onshore and offshore on the seabed, that have been identified as being of distinct relevance to paleoseismology. This increase in the number and variability of natural objects has been in response to the need for an improved and wider-ranging characterization of the seismic potential of inaccessible seismogenic sources, such as subduction zones and intracontinental seismic sources (e.g., New Madrid seismic zone). This has led to the identification of numerous indirect indicators and to the understanding of their functioning. Consequently, the appropriate paleoseismic approach is intimately related to the seismic landscape under assessment and is defined as the cumulative geomorphological and stratigraphic effect of the signs left on the environment of an area by its past earthquakes over a geologically recent time interval (Vittori et al., 1991; Michetti et al., 2005b). Likewise, the use of other natural paleoseismic markers has partly decreased, been abandoned, or fallen into disuse because of their limited applicability or limited results in terms of quantifying the seismic potential, for any or several of the following reasons: (1) The occurrence of the object is restricted to a specific region, a climate, or process, like lake varves, tree rings, or coral bands. (2) There is high specificity of the object to local conditions, e.g., scarp degradation to highly variable diffusion equations, lichenometry. (3) There is limited or unlikely preservation of the object in the geologic record, such as dead trees for dendrochronology. (4) There is difficulty in determining the magnitude and location of past earthquakes based on the regional distribution of the natural objects, which results in great uncertainty in assigning the observed seismically induced effects to a single common earthquake. (5) There is a degree of uncertainty in the cause-effect association between the natural indicator and earthquakes, for instance, liquefaction or mass movements. In the ultimate instance, the final selection of the natural indicators available for use is the responsibility of the paleoseismologist in accordance with those objects that each environment offers in terms of natural features. Finally, if available, direct objects are certainly everyone’s preference because they provide straightforward answers to the seismic history and potential for future rupture of a given fault or fault segment. ACKNOWLEDGMENTS We profusely thank Hans Diederix (private consultant, Colombia) and Pablo Silva (Universidad de Salamanca, Spain) for fruitful comments and suggestions on an earlier version of this contribution. Many thanks are also due to Marina Peña for her wonderful china-ink, hand-made drawings. As invited editors of this Geological Society of America Special Paper, we would like to acknowledge the help and work provided by James McCalpin during this book’s earlier stages. This volume is a

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MANUSCRIPT ACCEPTED BY THE SOCIETY 7 DECEMBER 2010

Printed in the USA

The Geological Society of America Special Paper 479 2011

Paleoseismicity of a low-slip-rate normal fault in the Rio Grande rift, USA: The Calabacillas fault, Albuquerque, New Mexico J.P. McCalpin GEO-HAZ Consulting, Box 837, Crestone, Colorado 81131, USA J.B.J. Harrison Department of Earth and Environmental Sciences, New Mexico Institute of Mining and Technology, Socorro, New Mexico 87801, USA G.W. Berger Desert Research Institute, 2215 Raggio Parkway, Reno, Nevada 89506-1095, USA H.C. Tobin Department of Geoscience, University of Wisconsin–Madison, Madison, Wisconsin 53706, USA

ABSTRACT The Calabacillas fault is a 40-km-long, down-to-the-east normal fault that trends N-S on the western edge of the Llano de Albuquerque, in western Albuquerque, New Mexico. It is one of several east-dipping normal faults that define the western margin of the Rio Grande rift at the latitude of Albuquerque. In the past 0.5–1 m.y., since the abandonment of the Llano de Albuquerque surface by the Rio Puerco and Rio Grande, vertical displacement on the Calabacillas fault has created a 27-m-high, eastfacing fault scarp on the western edge of the llano, equating to a long-term slip rate of 0.027–0.054 mm/yr. Our two trenches were located ~1 km from the south end of the fault, where a 1-km-wide graben has formed east of the main fault scarp. Trenching of the graben across the southern Calabacillas fault was 50% successful. The paleoearthquake event history on the 5.3-m-high antithetic scarp could not be reconstructed in detail because a strong carbonate soil profile had overprinted the entire 3-m-thick colluvial wedge deposit. It appears that numerous submeter displacements created this scarp, but the displacement was partitioned across several faults, so no single free face was higher than 10–20 cm. Free faces so small did not create colluvial wedges, and thus faulting did not trigger the pattern of footwall erosion and hanging-wall deposition needed to identify individual faulting events. On the 27-m-high main fault scarp, a 60-m-long trench straddled a minor slope break that overlies the main strand of the Calabacillas fault. The upper four soils exposed in the trench could be correlated across the main fault and indicated displacements of 10 cm, 30 cm, 55 cm, and 20 cm in the latest four paleoearthquakes.

McCalpin, J.P., Harrison, J.B.J., Berger, G.W., and Tobin, H.C., 2011, Paleoseismicity of a low-slip-rate normal fault in the Rio Grande rift, USA: The Calabacillas fault, Albuquerque, New Mexico, in Audemard M., F.A., Michetti, A.M., and McCalpin, J.P., eds., Geological Criteria for Evaluating Seismicity Revisited: Forty Years of Paleoseismic Investigations and the Natural Record of Past Earthquakes: Geological Society of America Special Paper 479, p. 23–46, doi:10.1130/2011.2479(01). For permission to copy, contact [email protected]. © 2011 The Geological Society of America. All rights reserved.

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McCalpin et al. Six infrared-stimulated luminescence (IRSL) dates on eolian sands range from 14 ka at a depth of 0.5 m to 219 ka at a depth of 5 m. Secondary calcium carbonate has accumulated in soils here at a rate of 0.17–0.35 g/k.y. The latest four faulting events are dated at ca. 14 ka, 23 ka, 35 ka, and 55 ka. Thus, the displacement and recurrence times increase with increasing age, yielding relatively consistent slip rates of 0.011–0.028 mm/yr. There is evidence at this trench for a late Pleistocene (14 ka) small faulting/cracking event, similar in displacement and timing to the youngest warping event interpreted for the County Dump fault, which lies ~5 km to the east. The displacements measured in the main scarp trench are even smaller than those inferred on the County Dump fault, despite the length of the Calabacillas fault (40 km) being similar to that of the County Dump fault (35 km). If our trenches had been located farther north toward the center of the Calabacillas fault, the displacements may have been larger. The ages and recurrence intervals of the four events that occurred subsequent to 55 ka are similar to those seen at the County Dump site. The youngest event on the Calabacillas fault had only 5–10 cm of throw, which is considerably smaller than the 20–55 cm throws of the three previous events. This situation parallels the County Dump chronology, where the youngest warping event was abnormally small compared to earlier events.

INTRODUCTION Purpose and Scope of Study This study is part of a continuing effort by the U.S. Geological Survey’s National Earthquake Hazard Reduction Program (NEHRP) to characterize the mid- to late Quaternary activity of normal faults in the Rio Grande rift near and in Albuquerque, New Mexico. This study continues efforts begun in 1996 (McCalpin et al., 2006) to trench faults that displace the surface of the Llano de Albuquerque (also known locally as West Mesa), an abandoned valley floor of the Rio Grande valley that now stands ~100–150 m above the Rio Grande on the western margin of metropolitan Albuquerque. The fault scarps that traverse the llano surface trend north-south and face east, ranging from 10 to 30 m high, but they have very gentle slope angles (typically less than 5°). During May 1999, we spent 3 weeks excavating, logging, and sampling two trenches on the southern end of the Calabacillas fault on the Llano de Albuquerque (West Mesa) of the Albuquerque Basin. The trenches were oriented east-west and transected north-south–trending fault scarps on either side of a north-south–trending, 1-km-wide graben (Fig. 1). The goal of this trenching investigation was to reconstruct the chronology of surface-faulting events on the Calabacillas fault since the abandonment of the Llano de Albuquerque as a depositional surface, in the past 0.5–1 m.y. The Albuquerque–Santa Fe area contains numerous late Quaternary normal faults, but it has not experienced a surfacerupturing earthquake in the 146 yr since the first recorded New Mexico earthquake (1849). However, there have been 10 historic earthquakes of Modified Mercalli Intensity V or greater in the urban corridor (Von Hake, 1975). The largest of these events

was the 18 May 1918 earthquake (MMI VII–VIII) at Cerillos (near Santa Fe), during which people were thrown off their feet and a break in the ground surface was noted (Von Hake, 1975). These earthquakes, plus the presence of Quaternary fault scarps, indicate the potential for larger, surface-rupturing earthquakes (M >6.5) in the corridor. Machette (1998, p. 89) stated: Because the level of seismicity of the Rio Grande rift is generally low, the populace (in general) believes that earthquakes do not pose a significant threat to them, whereas the presence of abundant young faults tells a different story.… A myopic view of earthquake hazards posed by individual structures (i.e., having recurrence intervals of 10,000 yr or more) can lead to a complacent attitude that strengthens a perception of low seismic potential for the region. Without proper caution, this attitude can be manifested in inappropriate construction styles, building codes, land-use policies, and the siting (or relocation) of important critical facilities.

Machette (1998) touches on three unique geologic situations that have led to the perception of low earthquake risk in Albuquerque: (1) long recurrence intervals on Quaternary faults, (2) the short, segmented nature of faults in the basin, and (3) the large number of faults. Situation 2 and trench studies described later herein suggest that most basin faults rupture in M 6.5–7 earthquakes near the threshold of surface rupture, resulting in 0.1–0.5 m of surface offset, rather than the >1 m observed on longer range-bounding faults. Situation 1 means that these small surface scarps are nearly obliterated by erosion in the long time interval between events. Situations 1 and 2 explain why there are hundreds of meters of cumulative vertical displacement of the Tertiary rift sediments (Santa Fe Group) on many intrabasin faults but fault scarps are unimpressive.

Paleoseismicity of a low-slip-rate normal fault in the Rio Grande rift, USA

Figure 1. Map showing faults with known and suspected displacements in Quaternary deposits near metropolitan Albuquerque, New Mexico (light gray). Dark-gray pattern shows pre-Neogene bedrock of the Sandia and Manzano Mountains. Fault numbers correspond to those in Table 1 of Personius et al. (1999). Faults with known displacements in the late Pleistocene (10–130 ka) or Holocene (59.1 ± 4.5

65.2 ± 5.6

30.2 ± 2.2

28.7 ± 2.4

24.9 ± 1.7

5.4 ± 0.4

5.5 ± 0.4

6.0 ± 0.4

5.3 ± 0.4

IRSL or OSL age (ka)**

Predates event V and buried soil, S1, on Llano de Manzano

Postdates event V

Postdates event V and predates event W

Postdates event W

Postdates event X and predates event Y(?)

Postdates event X and predates event Y(?)

Postdates event Z

Postdates event Z

Postdates event Z

Postdates event Z

Comments

CHSF02-4

UIC1053

Predates event V and S1 on Unit 1b 2.86 ± 0.38 7.21 ± 1.09 2.14 ± 0.02 315.00 ± 1.16 10 ± 3 3.60 ± 0.17 82.5 ± 6.0 Slope wash Llano de Manzano Note: Samples CHSF02-9 and CHSF02-11 were green optically stimulated luminescence (OSL) analyses, all others were infrared-stimulated luminescence (IRSL) analyses with measurements made using a multiple aliquot additive dose method, measuring blue emissions on the 4 to 11 µm polymineral fraction. *Sample locations are shown on Figures 5A and 5B. † De—equivalent dose. § All errors are at 1σ and were calculated by averaging the errors across the temperature range. **Ages are rounded to the nearest 100 yr, and all errors are at 1σ.

Unit 2 Sag pond deposit

Unit 4 Eolian sand

Unit 6 Loess

Unit 6 Loess

Unit 11 Scarp-derived slope-wash colluvium

Unit 13 Eolian sand

UIC1088

CHSF02-10

Unit 14 Eolian sand

UIC1357

Stratigraphic unit

TABLE 1. LUMINESCENCE DATA FOR THE CARRIZO SPRING TRENCH ACROSS SPLAY L OF THE HUBBELL SPRING FAULT SYSTEM

Lab sample no.

CHSF02-11

Field no.*

61

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Olig et al.

Unit 2 is only locally present on the downthrown side of FZ3 (Fig 5B). It pinches out to the east at st. 37 m and is faulted and warped down at its west end below the base of the trench west of st. 43.5 m. This dark-brown, sandy silty clay overlies unit 1b and is overlain by units 3a and 4. It is weakly bedded with fine sand partings, mottled with MnO staining, and generally lacks carbonate. Based on its location, limited extent, and stratigraphic characteristics, we interpret unit 2 to be a sag pond deposit at the base of a normal fault scarp created during event V by warping and offset on faults (FZ1, possibly FZ2, and FZ3 on Fig. 5B). Due to the subsequent dissolution of carbonate in unit 1b, stratigraphic relations are unclear as to whether unit 2 was deposited before or after development of S1. One IRSL sample from unit 2 (CHSF02-1 on Fig. 5B) yielded an age of 65.2 ± 5.6 ka, whereas another sample (CHSF02-2 on Fig. 5B) yielded a minimum age of >59.1 ± 4.5 ka due to saturation (Table 1). These ages are consistent with the IRSL ages from the underlying unit 1 and the degree of intervening soil development of S1, suggesting that S1 developed before event V occurred and before deposition of unit 2. Unit 3 is dominantly a pink fine sand that is faulted, warped down to the west, and appears to have been eroded away west of st. 38 m. We interpret unit 3 to be younger piedmont deposits of the Llano de Manzano, and similar to unit 1, unit 3 includes two subunits, 3a and 3b, where the latter is distinguished by a buried soil, S2. Subunits 3a and 3b are similar to and unconformably overlie units 1b and 1c, respectively. Unit 3a is a pinkish silty sand that contains some small gravel stringers and intraclasts. This unit is similar to the upper portion of unit 1b, and we interpret it be a mix of slope-wash and eolian deposits. Unit 3a is stratigraphically continuous with unit 3b, but it lacks the carbonate of the stage II nodular buried soil (S2) included in unit 3b. Similar to S1, the carbonate in S2 appears to have been dissolved from unit 3b west of st. 28 in the deformation zone. East of st. 18 m, the carbonate in S2 increasingly overprints the underlying S1 buried soil, and they became indistinguishable east of st. 11 m. Unit 4 unconformably overlies the buried S2 soil and units 3 and 2. Although it is warped and cut by fractures in the deformation zone, unit 4 did not appear offset by faults. This reddish, coarse to fine sand contains discontinuous coarser beds. It thickens on the downthrown side of faults, where sediments are generally coarser and include some intraclasts. Based on these characteristics, we interpret unit 4 to be colluvium and eolian sediments deposited after a faulting event, event W. However, similar to observations made by Personius et al. (2001) in the Hubbell Spring trench, this postfaulting unit is dominated by eolian deposition, and thus does not show many of the typical characteristics of fault-scarp–derived colluvial wedge deposits along normal faults, such as distinct debris and slope-wash facies forming wedged-shaped deposits (Nelson, 1992). An IRSL sample from unit 4 (CHSF02-6 on Fig. 5A) yielded on age of 30.2 ± 2.2 ka. This sample was collected from an eolian-dominated portion in the footwall because sediments in the deformation zone showed undesirable characteristics for luminescence dating (intraclasts, coarser sand, and extensive FeO and MnO staining).

Thus, although the eolian sediments that were sampled clearly postdate faulting, they may have been deposited well after faulting, and their age may not provide a close minimum limiting age for event W. Unit 5 is a well-sorted, reddish, fine to medium eolian sand that contains faint planar laminations, little carbonate, and extensive FeO and MnO staining. It conformably overlies unit 4 and includes a buried soil, S3, which apparently extends nearly the full length of the trench. However, the S3 catena varies considerably across the trench exposure (cf. soil profiles CS2, CS5, and CS1 in the Appendix), probably due to the different soil-forming conditions across the slope of the fault scarp. S3a extends from st. 0 to st. 8 m and is characterized by a mottled Btk horizon with a stage II to III carbonate morphology. Carbonate coatings on peds suggest overprinting of the original Bt horizon, likely similar to S3b, by carbonate related to development of overlying buried soils, S5 and possibly S6. S3 is apparently eroded away between st. 8 and st. 9 m. S3b extends from about st. 9 m to about st. 35 m and is characterized by a Bt horizon that is as much as 0.5 m thick and well cemented with sesquioxides. S3 is stripped away between st. 35 and st. 37. S3c extends from about st. 37 to the west end of the trench and is characterized by a mottled Btk horizon similar to S3a but more disseminated. S3c entirely overprints unit 5 and the top of unit 4, obscuring the contact between these units in the hanging wall. Although unit 5 is cut by fractures and warped down to the west with a total apparent throw of 2.8 ± 0.6 m down to the west (as measured in the trench and boreholes), we observed no discernible fault offsets of this unit. However, unit 5 thinned dramatically between st. 30 and 38 m. The top, including S3, is clearly stripped at the crest of the zone of maximum warping. The angular discordance between the base of unit 5 and the overlying stratigraphic and pedologic horizons suggests that this erosion occurred after a warping event, event X, and before unit 6 was deposited. Unit 6 unconformably overlies unit 5. Unit 6 is a buffcolored, silty, very fine eolian cover sand that is very well sorted and relatively homogeneous, lacks carbonate, and has some faint planar laminations. It appears to have been eroded away east of st. 9 m. West of st. 35.5 m, it is overprinted by carbonate from the overlying soil horizon, S5, and becomes indistinguishable from underlying and overlying units. Two IRSL samples from unit 6 (CHSF02-5 and CHSF02-7 on Figs. 5B and 5A, respectively) yielded an average age of 26.8 ± 2.4 ka (Table 1). Unit 7 is a pinkish tan, clayey silty fine sand that conformably overlies unit 6, extending from about st. 17 m to st. 32 m. This very well-sorted eolian cover sand is similar to unit 6, but it includes a weakly developed buried soil horizon, S4, that varies between a Bw and Btj horizon. Both units 6 and 7 are cut by fractures and warped, but we observed no discrete fault offsets in these units. In the deformation zone, both units 6 and 7 are overprinted, in angular discordance, by carbonate from the overlying buried soil horizon, S5, developed in unit 8. We believe this angular discordance was created by a small warping event, event Y(?),

Evidence for noncharacteristic ruptures of intrabasin faults in the Rio Grande rift because S5 appears to have formed on a slope that was steeper and higher than the slope where units 6 and 7 were deposited. Regardless, it is likely that unit 8 in the hanging wall includes units 6 and 7 near its base, but stratigraphic characteristics and contacts are obscured by soil formation in S5. Indeed, unit 8 is entirely characterized by a buried K horizon, S5, which obscures most of the original depositional characteristics of this pinkish-white, silty sand with clay (likely primarily an eolian deposit). It is present in the footwall and hanging wall but locally missing below the crest of the scarp, from about st. 17 m to 28 m, probably due to stripping. Although unit 8 is cut by fractures, warped down to the west, with a total apparent throw of 2.1 ± 0.5 m, and the top appears back tilted in the hanging wall and locally deformed adjacent to FZ4 (Fig. 5B), no faults with significant discrete slip were discernible in this unit. This is somewhat surprising because the abrupt upper contact with the overlying unit 9 makes a relatively distinct marker that is deformed but does not appear offset, even though unit 10 above is clearly offset by the upper portion of FZ4 as it steps westward. It appears that deformation in this portion of FZ4 is very distributed, and discrete slip is discontinuous. Soil carbonate accumulation may also mask small offsets. Unit 8 varies from a stage II to III+ carbonate morphology, but it is dominantly a stage III. It is also more than double in thickness in the hanging wall (over 1 m), where it appears more disseminated than in the footwall and contains large zones that are punky and friable in texture. It also contains irregular-shaped carbonate nodules, some as large as cobbles, that weather out of the trench wall easily in places due to apparent partial dissolution of carbonate in the surrounding matrix. We believe unit 8 has been extensively affected by groundwater and/or spring water upwelling along the faults and fractures in the deformation zone. This would not only explain the unusual textures, but also the apparently incongruous large amount of carbonate accumulation in a horizon that is younger than 30 ka, as indicated by luminescence ages from unit 6. Unit 9 is a silty, clayey sand that is locally present only on the downthrown side of the deformation zone between st. 32 and 56 m. It is characterized by a buried soil, S6, which consists of a mottled, friable Btk horizon with a stage II to III carbonate morphology. However, similar to unit 8, we believe the carbonate in unit 9 may not all be pedogenic, and thus the morphology is not a reliable indicator of age. Unit 9 contains reworked carbonate nodules, apparently from unit 8, which generally decrease downslope. Based on these characteristics and its limited extent on the downthrown side of the deformation zone, we interpret unit 9 to have been scarp-derived colluvium and eolian material deposited after the warping event Y(?). Unfortunately, due to the extensive carbonate accumulation and soil development, unit 9 was not a good candidate for luminescence age analyses. Unit 9 was also cut by faults of FZ4b during the most recent event, event Z. Additionally, unit 9 shows weak planar laminations in places, and these are tilted and warped on the downthrown side of FZ4. Unit 10 is a silty sand with clay that is characterized by the youngest buried soil in the trench, S7. This soil is a stage II nodu-

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lar carbonate horizon that extends the full length of the trench, except where it was broken up by faults of FZ4b during event Z and disturbed by krotovina between st. 39.5 and 41.5 m. Additionally, unit 10 appears slightly back tilted toward the east on the downthrown side of FZ4b. Similar to unit 9, the carbonate accumulation and soil development in unit 10 prevented luminescence age analyses. Unit 11 is a light-brown, silty sand that contains blocks of unit 10 and reworked carbonate nodules. It is a wedge-shaped deposit that unconformably overlies the buried soil on unit 10. It is not offset by any faults, and it is present only on the downthrown side of FZ4b. Based on these characteristics, we interpret unit 11 to be primarily fault-scarp–derived colluvium with minor eolian material deposited after the youngest faulting event, event Z. Two luminescence samples (CHSF02-8 and CHSF02-9 on Fig. 5B) yielded an average age of 5.5 ± 0.4 ka for the distal and eolian-dominated portion of unit 11. Units 12 and 13 are, respectively, silty and coarse sand deposits that are dominantly eolian sediments with some slope wash. These lenticular-shaped deposits are cut by open fractures, but they did not appear faulted or warped. They conformably overlie unit 11 and appear to be filling the topographic depression that was likely created during faulting event Z, although it is possible that deposition of units 12 and 13 was related to another younger, minor warping event that could have occurred subsequent to event Z and deposition of unit 11. However, given the lack of any other evidence for this hypothetical event, we believe it is much more likely that the open fractures were related to nontectonic processes (such as settlement, bioturbation, freezethaw, etc.) and that the deposition of distinct eolian packages was related to nontectonic causes, such as climate change. An IRSL sample (CHSF 02-10 on Fig. 5B) yielded an age of 6.0 ± 0.4 ka, which is stratigraphically consistent (within 1σ error) with the age for the underlying unit 11 and the overall lack of soil development in units 11 through 13. Finally, unit 14 is a loose, tan, fine sand that drapes the scarp. This eolian sand includes the modern soil, which is characterized by roots in a weakly developed BW horizon. A luminescence sample (CHSF02-11 on Fig. 5B) collected from below the modern soil in unit 14 yielded a luminescence age of 5.3 ± 0.4 ka (Table 1). Trench Structure The broad deformation zone exposed in the trench is characterized by: (1) a zone of warping down to the west, which is coincident with the scarp face; (2) a narrower zone of near-vertical and bedding-parallel fractures, between st. 28 m and st. 41 m, which does not show discernible discrete slip (shown in black on Figs. 5A and 5B); and (3) a series of four west-dipping to subvertical fault zones (FZ1–FZ4, shown in red on Fig. 5B) that offset strata down to the west and are roughly coincident with the zone of fractures, the maximum zone of warping, and the maximum slope angle on the scarp. Although warping appears to have deformed units 1 through 10 in a broad zone (i.e., from the crest of the scarp westward),

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there is a more concentrated zone of warping between st. 28 and 42 m that is coincident with faulting and the maximum scarp angle. Within this zone, older units are generally warped more than younger units, but differential warping and distinct events are much more difficult to distinguish than differential offsets and events on faults. However, a progression of warping (and faulting) events is still evident from the overall thickening of many units and soils in the hanging wall, the erosion of several units and soils in the zone of maximum deformation, and the angular discordances between some units and overlying units and soils, as discussed in the previous section. Based on these relations and total differential offsets on buried soils (S1, S2, S3, S5, and S7) (discussed in “Deformation Event Summary and Chronology” section), it is evident that warping occurred during all of the deformation events on splay L of the Hubbell Spring fault system, albeit to various degrees. Fault terminations at the top of buried soils and differential offsets on specific faults provide evidence for three of the faulting events, events V, W, and Z. Fault-zone orientations and dip-slip measurements are shown on the trench log (Fig. 5B). Faults FZ1 and FZ2 both cut units 1 and 3, terminating at the top of unit 3a, the stratigraphic equivalent to the top of the buried soil S2 and the event horizon for event W. Differential offset between units 1 and 3 suggest that FZ1 and FZ2 were both active during events V and W. However, given the uncertainties in measurement, larger offsets of unit 1 may be partially due to the listric geometry of FZ1. Similar to FZ1 and FZ2, FZ3 cut units 1 and 3, terminating at the top of unit 3, although fractures (without offset) extended into unit 4. Small differential offsets between units 1 and 3 were apparent but not definitive on FZ3, given the warping and complex fault geometry of this zone. FZ4 is even more complex. It is characterized by a wide distributed shear zone located between st. 39 and 41 m (Fig. 5B), with discontinuous fault strands that step to the west stratigraphically upward and include lower (FZ4a) and upper (FZ4b) portions. The lower portion, FZ4a, cuts units 1 and 2 and the base of unit 4 and has many associated fractures that extend upward into unit 5. Although the upper portion, FZ4b, offsets units 9 and 10 with discrete slip on faults, definitive discrete slip offset of underlying units 8 and 5 is not apparent, although these units are clearly warped and deformed down to the west between FZ4a and FZ4b. It would appear that shearing is so distributed in this part of FZ4 as to be visible only as warping, although relations are also somewhat complicated by carbonate accumulation and overprinting of S3 and S5 soil development, which may obscure some discrete slip. The upward termination of FZ4b is at the top of the buried soil S7 in unit 10, the event horizon for event Z, with open fractures extending into overlying units 11 and 12. Differential offsets of unit 10 versus unit 3b indicate that FZ4 was active during events W and Z. Soil Pit Exposure The soil pit exposed an extensively bioturbated package of dominantly eolian sand with minor alluvium that included three buried soils underlying the modern soil on loose eolian sand

(Fig. 6). The lowermost soil is a disseminated K horizon developed on a sandy silt to silty sand with gravel. It contains the most carbonate (stage II to III) and lacks the clay, mottling, and peds that are characteristic of S3 in the trench. Therefore, we think this soil most likely correlates to S5 in the trench. Overlying the K horizon, there is a mottled, silty very fine sand with a buried disseminated Bk horizon. This soil may correlate to S6 in unit 9 or S7 in unit 10 in the trench, although carbonate accumulation appears much more disseminated in the Bk horizon. Overlying the Bk horizon, there is sandy alluvium with a thin Btjk horizon that does not appear to directly correlate to any soils in the trench, although it may be time-equivalent to S7. The sandy alluvium appears to have been deposited by a small drainage that incises the fault scarp north of the trench. Due to the extensive bioturbation and more disseminated character of the buried soils exposed in the soil pit, correlations to soils in the trench were slightly ambiguous. However, it is clear that neither of the buried soils, S2 or S1, nor the channel gravels of unit 1 were exposed in the soil pit, and presumably these soils and deposits are at a greater depth below the pit exposure. Indeed, as discussed in the next section, S1 and S2 were observed at greater depths in the boreholes. Boreholes The locations of the three drill holes, projected onto the scarp profile, along with the interpreted depths of buried soils S1 and S2, are shown on Figure 5D. Logs of borings are included in Appendix C of Olig et al. (2004). Boring B1 was located just south of st. 60 m at the west end of the trench. It was the deepest hole (total depth 10.39 m) and provided relatively good correlations to stratigraphy exposed in the trench, except for some uncertainties due to sloughing at depths around 1.5 and 2.5 m. At a depth of 5.6 m, boring B1 encountered a buried carbonate soil horizon that was mottled and appeared nodular. It was developed on a gravelly sand that coarsened downward to a clean sandy gravel with clasts of gneiss, quartzite, and other dark metamorphic rocks. Based on these characteristics, we correlate these sandy gravels to the channel deposits of unit 1a in the trench, and the soil to S1 in the trench. At a shallower depth of 3.85 m, boring B1 encountered another mottled carbonate soil horizon that appeared nodular. Our preferred interpretation is that this soil correlates to S2 in the trench. Alternatively, a thin, more clayey mottled carbonate horizon encountered at 4.5 m depth possibly correlates to S2. Boring B2 was located 46 m southwest of B1 and was 7.22 m deep. B2 encountered a mottled carbonate soil horizon at 2.53 m depth, which we believe correlates to S2 in the trench. Underlying that was another mottled carbonate horizon at 4.76 m depth, which we correlate to S1 in the trench. Underlying S1 in both B1 and B2 was a thick (>1 m), pinkish, well-sorted, silty fine sand. Borehole B3 was located between B1 and B2. Samples from the upper 2 m of B3 were disturbed due to drilling problems. However, a carbonate soil horizon developed on a sandy gravel with metamorphic clasts was encountered at a depth of

Evidence for noncharacteristic ruptures of intrabasin faults in the Rio Grande rift 4.95 m, and we correlate this soil and gravel to S1 and unit 1 in the trench. At a shallower depth of 3.7 m, another carbonate soil horizon was encountered in B3, which we tentatively correlate to S2 in the trench. Unfortunately, at the bottom of B3, we could not punch through unit 1 and presumably into the underlying pinkish silty fine sand that was encountered in boreholes B1 and B2. Deformation Event Summary and Chronology This section summarizes the evidence for, and timing constraints on, the faulting and warping events at the Carrizo Spring trench site. Differential offsets provide important evidence for identifying events and information about event size. Due to the extensive warping at the Carrizo Spring site, we knew that we needed to look beyond just differential offsets on individual faults and compare the total down-to-the-west throw (including fault slip and warping) between event horizons. Using observations from both the trench and boreholes, we projected key buried soil horizons across the entire deformation zone to estimate the apparent throw on these key horizons. In particular, we were able to measure the apparent total throw on key marker horizons, buried soils S1, S2, S3, S5, and S7. Independent stratigraphic and structural evidence, which is summarized in the following sections, indicates or suggests that the tops of these buried soils correspond to the event horizons for paleoearthquakes V, W, X, Y(?), and Z, respectively. We refer to these total throw measurements as “apparent” because our correlations of units across the deformation zone rely heavily on pedogenic characteristics, and soils can form on preexisting slopes such that all of the vertical relief we measured may not be related to offset that occurred after the soil formed. However, the generally subrounded to rounded gravels of unit 1a, with clasts as large as boulders, were deposited by relatively high-energy streams flowing west to southwest from the Manzano Mountains, out across the Llano de Manzano surface, and across any scarp associated with splay L. Given the nature of the channel gravels comprising unit 1a, we believe it is likely that these streams beveled off any preexisting scarp before soil S1 formed. Therefore, in our analysis, we assume that S1 formed on a relatively uniform and gentle slope without a preexisting scarp. Assuming there was relatively little to no initial relief on unit 1, and because the vertical relief measured on each event horizon is progressively larger for each older event, we make the key inference that all this post–unit 1 vertical relief was tectonically created in a series of successive faulting/deformation events. This seems reasonable because we see no other likely cause for repeated creation of new relief on a north-trending scarp at this location, especially given the nature of all the deposits overlying unit 1, which are all generally finegrained eolian and slope-wash sediments. Table 2 summarizes the total apparent throw measured on the tops of the buried soils (S1, S2, S3, S5, and S7) respectively forming the event horizons for events V, W, X, Y(?), and Z. These can be used to calculate differential displacements and estimate the displacements per event as follows. Assuming S1 formed on a

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gentle slope with no scarp as previously discussed, we measured 7.3 ± 1.0 m of total cumulative throw on this buried soil by projecting it across the deformation zone from the trench exposures and boreholes (Fig. 5D; note that we actually measured displacements on a 1:40 scale version of Fig. 5D that was significantly reduced here for publication). Similarly, from the trench exposure and boreholes, we measured 6.5 ± 0.3 m of total apparent throw on soil S2. We calculated a differential offset between S1and S2 of 0.8 m to estimate the net vertical tectonic displacement for event V (Table 2). The net vertical tectonic displacement is the measure of vertical slip at the fault that accounts for the back tilting and antithetic faulting that is common along normal faults (Swan et al., 1980). Although some minor back tilting was apparent for a couple of events, more importantly, at the Carrizo Spring site, this measure also includes the effects of warping. We similarly used differential offsets between each successive event horizon (S2, S3, S5, and S7) to estimate net vertical tectonic displacement for each subsequent deformation event. Resulting preferred estimates of net vertical tectonic displacement per event range from 0.4 to 3.7 m (Table 2), indicating highly variable and noncharacteristic behavior (discussed further in “Rupture History and Behavior” section). The displacements are smallest for the two warping events, X and Y(?). Indeed, due to the small preferred offset for event Y(?), and because independent stratigraphic and pedologic evidence was not as strong for this event, we recognize the uncertainty in the occurrence of this probable warping event by explicitly using a query in referring to it. Except for the youngest event, Z, uncertainties for displacements per event are large due to the difficulties in projecting horizons across the broad deformation zone, and the cumulative effect of calculating differential offsets. This results in a minimum estimate of 0 m for three of these events (V, X, and Y[?]). However, with the possible exception of event Y(?), we believe that the independent evidence for the occurrence of events V and X argues for some minimum offset larger than 0 m for these events. It is also worth pointing out here that the cumulative throw on S1 is 7.3 ± 1.0 m, which is similar to the 7.5 ± 1.0 m of net vertical tectonic displacement measured across the Llano de Manzano surface on the scarp profile. At first glance, this seems somewhat surprising because offsets measured from scarp profiles on normal faults might be considered minimums due to expected postfaulting erosion of the footwall and deposition in the hanging wall. However, the stratigraphic record preserved in the Carrizo Spring trench shows little evidence for preferential erosion of the footwall aside from at the crest of the deformation zone. Additionally, although fault scarps provide excellent local sediment traps for eolian sands in the region (e.g., Personius and Mahan, 2003; McCalpin et al., 2006), deposits can also blanket scarps, so that sediment is not only accumulating in the hanging wall but also in the footwall (e.g., units 4, 5, 6, and 7). These types of cover sands and intervening buried soils are likely present elsewhere on the Llano de Manzano surface and would provide excellent stratigraphic markers for paleoseismic studies

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Olig et al. TABLE 2. ESTIMATES OF DISPLACEMENTS ON SPLAY L OF THE HUBBELL SPRING FAULT SYSTEM

Deformation event

Event horizon

Total apparent throw on event horizon* (m)

Estimated net vertical tectonic displacement per event (m) 0.8 † (>0 , 1.6)

Basis for estimate

Event V

Top of buried soil S1 (on unit 1)

7.3 ± 0.3

Event W

Top of buried soil S2 (on unit 3)

6.5 ± 0.5

3.7 (2.6, 4.8)

Differential offset between total apparent throw on top of S2 and top of S3 (6.5–2.8 m)

Event X

Top of buried soil S3 (on unit 5)

2.8 ± 0.6

0.7 † (>0 , 1.8)

Event Y(?)

Top of buried soil S5 (on unit 8)

2.1 ± 0.5

0.4 (0, 1.2)

Event Z

Top of buried soil S7 (on unit 10)

1.7 ± 0.3

1.7 (1.4, 2.0)

Differential offset between total apparent throw on top of S3 and top of S5 (2.8–2.1 m) Differential offset between total apparent throw on top of S5 and top of S7 (2.1–1.7 m) Total apparent throw on top of S7

Differential offset between apparent throw on top of S1 and top of S2 (7.3–6.5 m)

Style of deformation

Normal slip on FZ1, FZ2, and possibly FZ3 (and buried faults west of st. 43 m?) Warping Fracturing Normal slip on FZ1 through FZ4 (and buried faults west of st. 43 m?) Warping Fracturing Warping Fracturing

Warping Fracturing Back tilting (?) Normal slip on FZ4 Warping Fracturing (?) Back tilting (in FZ4 hanging wall)

Note: All displacements are down to the west. *Measured from trench exposures (and in boreholes for S1 and S2) by projecting horizons across the zone of deformation. Assumes channel gravels of unit 1 leveled off any preexisting scarp, and all subsequent vertical relief was created by tectonic deformation (see text for discussion). † See text for discussion of minimum values.

of other faults scarps. The drawback of these deposits is that they mute the geomorphic expression and can even bury faults, making fault scarps appear much more discontinuous and difficult to map (Olig and Zachariasen, 2010). In retrospect, the geomorphic expression of fault splay L of the Hubbell Spring fault system at the Carrizo Spring trench site is consistent with the paleoseismic record exposed in the trench, with a very broad scarp resulting from a broad deformation zone of faulting and warping. Additionally, the maximum scarp angle was located at the zone of most concentrated faulting and warping. Scarp heights appear subdued due to little back tilting and a “bulge” of eolian deposits at the scarp base. The low maximum scarp angle and single bevel on the scarp profile may at first seem inconsistent with the four to five late Pleistocene events interpreted in the trench, but the single bevel is actually consistent with the dominant deformation style of warping, and the small maximum slope angle of 12° is likely more related to the angle of repose for deposition of eolian sand on the scarp rather than degradation of an older fault scarp. Thus, the use of morphometric comparisons of scarp height and maximum slope angle (e.g., Bucknam and Anderson, 1979) does not appear to be a reliable indicator of fault-scarp age in this environment, which explains why Machette and McGimsey (1983) erroneously concluded that

the youngest faulting was much older than 15 ka based on their analysis of scarps of the Hubbell Spring fault system. However, our observations also suggest that scarp profiles can still provide useful slip estimates despite the extensive eolian deposition. A summary discussion of the structural, stratigraphic, and pedologic evidence for each surface-faulting/deformation event and the timing of events are presented next. Event Z Structural, stratigraphic, and pedologic evidence for this youngest surface-faulting event includes: (1) offset (including discrete slip on fault FZ4b and warping) of the top of unit 10, and the buried soil S7, 1.7 ± 0.3 m down to the west; (2) fault terminations at the top of S7; (3) an overlying unfaulted, colluvial-wedge deposit (unit 11) adjacent to the maximum zone of faulting, which contained reworked carbonated nodules from unit 10, blocks of unit 10, as well as eolian sediment banked against the scarp created by event Z; and (4) slight back tilting of unit 10 to the east toward fault FZ4b. Event Z occurred sometime before deposition of the distal portion of unit 11, ca. 5–6 ka, and well after unit 6 was deposited at 24–29 ka, because not only were units 7 through 10 deposited subsequently, but three intervening buried soils (S4, S5, and S6) formed. However, as previously discussed, some of

Evidence for noncharacteristic ruptures of intrabasin faults in the Rio Grande rift the carbonate accumulation observed in S5 and S6 may be related to nonpedogenic processes, and so unfortunately the carbonate morphology of these soils is likely not reliable a indicator of their age. Regardless, the general lack of soil development in units 11–14 also suggests that event Z occurred closer to 6 ka than 24 ka and was likely younger than 15 ka. Event Y(?) The evidence for the penultimate event, event Y(?), is suggestive, but not conclusive because the apparent deformation associated with this event is limited to minor warping of unit 8 and the buried soil S5 ~0.4 m down to the west. No discrete slip on any faults was observed. However, the angular relation between the buried soil S5 and the underlying units 6 and 7 suggests that this soil formed on a slope that was steeper than the one units 6 and 7 were originally deposited on. This, and the apparent increased differential offset between S7 and S5 suggest that this deformation event occurred after S5 formed but before unit 9 was deposited. Indeed, unit 9 appears to be a mix of scarp-derived colluvium (including reworked carbonate nodules) and eolian sediment banked up against the scarp after event Y(?) occurred. The timing for event Y(?) is very poorly constrained. The Btk horizon of S6 formed on unit 9 suggests that event Y(?) was somewhat older than Holocene, and of course it is younger than deposition of unit 6 at ca. 24–29 ka. Event X Although event X did not result in any apparent discrete slip on faults, fracturing and warping of units 1–5 did result in 0.7 m of overall differential down-to-the-west net vertical tectonic displacement between the tops of buried soils S3 and S5 (see Appendix). Additionally, stripping of S3 and the top of unit 5 locally at the crest of the zone of maximum warping, along with the angular discordance between the base of unit 5 and the overlying stratigraphic and pedologic horizons, also indicates that erosion occurred after a warping event of unit 5 and S3 but before deposition of unit 6. Based on luminescence ages for units 4 and 6 (Table 1), event X occurred between 26.8 ± 2.4 ka and 30.2 ± 2.4 ka. Note that if event Y(?) did not occur, it would imply that larger displacements occurred during event X, ~1.1 m instead of only 0.7 m. Event W Compelling stratigraphic and structural evidence for event W includes: (1) discrete offset of units 1–3 along faults FZ1–FZ4, with multiple fault terminations at the base of unit 4 (the stratigraphic equivalent to the top of the buried soil S2); (2) warping and faulting resulting in large differential down-to-the-west offset of 3.7 m between the top of buried soils S2 and S3 (Appendix); and (3) thickening and coarsening of unit 4 on the downthrown side of faults FZ1, FZ2, and FZ3. Despite the large net vertical tectonic displacement indicated for event W, displacements on individual faults are small and distributed across multiple splays, forming wide zones. Additionally, similar to other events, warp-

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ing is also evident across a wide zone. One somewhat anomalous stratigraphic relation is the apparent thinning of unit 4 west of st. 38 m, on the downthrown side of FZ4. This may be due to the presence of buried faults west of st. 43 that were not exposed in the trench. Alternatively, it may be due to stratigraphic uncertainties, because the contact between units 4 and 5 becomes more diffuse and affected by overprinting of soil carbonate west of st. 37 m. Regardless, our estimated value of 3.7 m of net vertical tectonic displacement for event W based on drill-hole and trench data includes any offset on those potential buried faults. The timing of event W is not tightly constrained, but it occurred sometime before deposition of unit 4 (ca. 30.2 ± 2.2 ka) but well after deposition of unit 2 (65 ± 6 ka), and after subsequent deposition of unit 3 and formation of the stage II carbonate horizon of S2 on unit 3. Event V Stratigraphic, structural, and pedologic evidence from the trench and boreholes indicates that the oldest event, event V, resulted in ~0.8 m of net vertical tectonic displacement of the top of S1, the buried soil developed on unit 1. Deformation is characterized by warping, fracturing, and discrete normal slip on FZ1, FZ2, and possibly FZ3. In addition to differential offsets, a slight thickening of unit 3a on the downthrown side of faults FZ1 and FZ2 also supports minor slip on these faults during event V. However, much of the deformation associated with event V appears to have been accommodated by warping of unit 1 and S1, creating a depression at the base of the scarp, where a sag pond deposited unit 2. The timing of event V is constrained to be well after unit 1 was deposited (ca. 84 ± 6 ka) and the S1 soil subsequently formed but likely shortly before deposition of unit 2 ca. 65.2 ± 5.6 ka. ANALYSIS OF PALEOSEISMIC PARAMETERS AND DISCUSSION OF FAULT BEHAVIOR Rupture History and Behavior The paleoseismic record of surface faulting and warping events that we deciphered for fault splay L of the Hubbell Spring fault system at the Carrizo Spring site is summarized in Figure 7, including timing constraints and preferred net vertical tectonic displacements per event. We found stratigraphic, structural, and pedologic evidence that indicates at least four, possibly five, surface-deforming earthquakes (events V through Z) occurred since a stage II buried soil carbonate horizon formed on sediments that were deposited on the Llano de Manzano ca. 84 ± 6 ka. Also summarized on Figure 7 is the record of surface-faulting events deciphered by Personius et al. (2001) and Personius and Mahan (2003) for the western Hubbell Spring fault (splay J of the Hubbell Spring fault system) at the Hubbell Spring site. They found evidence for four surface-faulting events that occurred after carbonate rinds formed on fan gravels ca. 92 ± 7 ka. The available paleoseismic data for the Hubbell Spring fault system suggests complex rupture behavior that includes both

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Splay L (Central) - HSFS

Splay J (Western) - HSFS s

k.y.

k.y.

k.y.

s

slope-wash

ka slope-wash

Figure 7. Comparison of the paleoseismic records of fault splay L (this study) and fault splay J (Personius et al., 2001; Personius and Mahan, 2003) of the Hubbell Spring fault system (HSFS).

Evidence for noncharacteristic ruptures of intrabasin faults in the Rio Grande rift independent rupture of splay L of the Hubbell Spring fault system and simultaneous rupture of splays L and J of the Hubbell Spring fault system. Although the timing constraints are poor for the first event on splay J and for events Y(?) and W on splay L, comparison of the paleoseismic records (Fig. 7) indicates that the timing of the four largest events on splay L (events Z, X, W, and V) overlaps with the timing of the past four events on splay J (fourth through first events, respectively), suggesting simultaneous rupture of splays L and J together during larger events. The relatively large displacements for these events (1–2 m on fault splay J and 0.7–3.7 m on fault splay L), and the similar amounts of total throw since the oldest event (5–8 m on splay J vs. 7.3 ± 0.5 m on splay L) also support simultaneous rupture of the two splays. Additionally, the buried soils that formed before each surface-deforming event appear to correlate between sites (e.g., our soil S1 correlates to the buried soil in unit 2 at the Hubbell Spring site, etc.; Fig. 7), which also supports simultaneous rupture. However, given the resolution of ages, even for the events with better age constraints (e.g., event X on fault splay L and the third event on fault splay J), we acknowledge that we cannot preclude the possibility that the events on each splay may have occurred separately. Regardless, the smallest event on fault splay L, event Y(?), does not appear to correlate to any events on fault splay J of the Hubbell Spring fault system. Assuming that this event did indeed occur and that the record deciphered for fault splay J is complete, this indicates that fault splay L also does occasionally rupture independently from fault splay J of the Hubbell Spring fault system in smaller events. Further comparison of the two sites also reveals additional similarities and differences in the paleoseismic records that warrant discussion. Some of the more significant stratigraphic and structural similarities between the sites are (1) the domination of eolian sedimentation along the scarps; (2) the wide distributed deformation zone of many faults across a single scarp; and (3) the box-like network of slope-parallel and subvertical carbonate-filled veins within the deformation zones. In regard to the first similarity, this study adds to a growing body of evidence (e.g., McCalpin et al., 2006; Personius and Mahan, 2003) for a stratigraphic signature along normal-slip intrabasin faults in the Rio Grande rift that is different than typical range-bounding normal faults. That is, deposits along intrabasin rift faults generally: (1) lack the debris-facies colluvial deposits that are typically proximal to scarps of normal, range-bounding faults; (2) are dominated by eolian sediments banking up against the scarp rather than colluvial sediments derived from the scarp; and (3) are more lenticular-shaped and can extend completely across the scarp as compared to the classic triangular, colluvial wedge deposit that is primarily limited to the downthrown side of the fault. The success of luminescence dating of the eolian and colluvial sediments at both Hubbell Spring fault system trench sites is promising for applications elsewhere in the Albuquerque Basin, not only for fault studies, but for all types of Quaternary studies, particularly where additional absolute age constraints are sorely needed to understand complex diachronous surfaces.

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There are also important stratigraphic, pedologic, and structural differences between the two Hubbell Spring fault system trench sites. One important difference between the sites is the probable occurrence of an additional event (event Y?) on fault splay L, and the more complex stratigraphic and pedologic sequence that predates and postdates this small warping event at the Carrizo Spring site. This highlights many questions about the seismogenic relation among all the fault splays within the Hubbell Spring fault system, and the need for additional paleoseismic studies to understand the behavior of the many other unstudied splays. Given the anastomosing geometry and close proximity between some splays of the Hubbell Spring fault system, it is likely that some splays merge together in the subsurface. However, the geometry of this complex system (its long length, great width, and the fact that nearly all the splays are west-dipping) also indicates that all of the fault splays of the Hubbell Spring fault system cannot merge together. Given the complex geometry and the available paleoseismic data, it seems unlikely that the entire Hubbell Spring fault system ruptures simultaneously, especially in every event. Still, preliminary observations of alongstrike displacement variations and the anastomosing geometry of faults do suggest complex patterns of slip transfer between faults, with northwest- and northeast-striking sections serving as relay ramps between principal north-striking faults (Olig and Zachariasen, 2010). Obviously, we need additional coordinated studies to better understand rupture patterns for the entire fault system and to more accurately model rupture behavior in hazard evaluations, which presently use very simplistic rupture models for the Hubbell Spring fault system due to the lack of data (e.g., Wong et al., 2004; Olig et al., 2007). Another structural difference between the trench sites is the greater degree of warping and variability in the displacements per event at the Carrizo Spring site. Displacements per event on fault splay L showed large variability, ranging from 0.4 to 3.7 m (Fig. 7). In contrast, estimated displacements per event for fault splay J only ranged between 1 and 2 m (Personius and Mahan, 2003); however, displacements for this splay are not as well constrained as for splay L. Despite the apparently more uniform displacements per event on splay J, cumulative displacements per event (assuming events were simultaneous in both trench sites) still show large variability. For the youngest (event Z) to the oldest (event V) events, we estimate cumulative displacements of 3.7, 0.4, 1.7, 4.7, and 2.8 m. Calculations of cumulative displacements are complicated by the fact that the two trench sites are over 18 km apart along strike, and displacements on a fault can vary significantly along strike. Here, we assumed that displacements measured at each site are representative averages for each fault splay, and so we simply added displacements per event from each site to estimate cumulative values for splays L and J. Although we believe that the complex rupture behavior interactions between different splays of this intrabasin fault are a significant contributing factor to the large variability in displacements per event, it is worth pointing out that displacement variability is still large between simultaneous rupture events, ranging from 1.7 to 4.7 m.

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The large variability in displacements per event indicates noncharacteristic behavior for splay L of the Hubbell Spring fault system, and likely for the Hubbell Spring fault system overall, which has important implications for recurrence models used in probabilistic hazard analyses (Wong et al., 2004). As originally defined by Schwartz and Coppersmith (1984), the characteristic recurrence model was based on paleoseismic observations along the San Andreas and Wasatch faults of similar-sized displacements per event at a particular point along a fault. The characteristic model predicts fewer moderate-size events and generally results in lower hazard estimates than the traditional Gutenberg-Richter exponential frequency-magnitude relationship (Youngs and Coppersmith, 1985). Thus, the large variability in displacements for the Hubbell Spring fault system implies noncharacteristic behavior and higher associated hazard. However, we also recognize that it is possible that along-strike variations in displacement on fault splay J are such that a trench on that splay at the same latitude as the Carrizo Spring site would reveal complementary displacements per event that would result in total displacements per event for both splays that were more similar in size. However, observed variations between displacements per event are so large that noncharacteristic behavior is still strongly suggested, regardless of possible along-strike displacement variations. For example, 2.6 m is an absolute minimum estimate for event W (see Table 2, and assuming the second event shows little or no displacement on fault splay J at the latitude of the Carrizo Spring site). In contrast, assuming event Y(?) did not occur on fault splay

J, 1.2 m is a maximum estimate for this event (Table 2), which still implies noncharacteristic behavior for the fault zone overall. Paleomagnitudes Paleomagnitude estimations for prehistoric deformation events on the Hubbell Spring fault system are complicated by the geometry of its many overlapping fault splays. Although many empirical relations have been developed to estimate paleomagnitudes from various fault parameters, such as length and displacement per event, none of these has been developed for faults with multiple, subparallel overlapping splays like the Hubbell Spring fault system. In particular, for simultaneous rupture of splays, it is not clear whether summation of displacements per event for each splay is appropriate or not, although we did add them here because it seems the most logical approach. Table 3 shows paleomagnitude estimates for the Hubbell Spring fault system using various relations based on surface-rupture length (L), average displacement (AD), and maximum displacement (MD). Paleomagnitude estimates vary from Mw 6.6 to 7.5 depending on the assumed input parameters and rupture behavior of the Hubbell Spring fault system. All of the estimates based on displacement could arguably be considered minimums because unknown possible displacements on other splays (other than J and L) have not been included. Additional investigation of other fault splays and a more complete paleoseismic record are obviously needed for the Hubbell Spring fault system to better estimate paleomagnitudes.

TABLE 3. PALEOMAGNITUDE ESTIMATES FOR SURFACE-FAULTING EARTHQUAKES ON THE HUBBELL SPRING FAULT SYSTEM Fault parameter L = 74 km*

Expected moment magnitude (Mw)* 7.25 ± 0.28

MD = 4.7 m

†

7.19 ± 0.39

AD = 1.7 m

§

7.12 ± 0.39

AD = 4.7 m

#

7.48 ± 0.39

AD = 0.4 m** ††

6.60 ± 0.39

L = 42 km 6.96 ± 0.28 Note: We used empirical relations from Wells and Coppersmith (1994) for all types of slip: Mw = 5.08 + 1.16(log L), σ = 0.28; Mw = 6.93 + 0.82(log AD), σ = 0.39; Mw = 6.69 + 0.74(log MD), σ = 0.04. All surface rupture lengths (L) were measured straight line, end to end in kilometers. MD is maximum displacement along strike, and AD is average displacement along strike per event measured in meters. *For the entire Hubbell Spring fault system and based on mapping by Olig and Zachariasen (2008). † This is the maximum observed displacement for simultaneous rupture of fault splays J and L, which occurred during event W; here assumed to be representative of the maximum along-strike displacement. § This is the minimum observed displacement of simultaneous rupture of fault splays J and L, which occurred during event X; here assumed to provide a lower bound for the average displacement for these splays. # Maximum observed displacement for simultaneous rupture of splays L and J, but here assumed to provide an upper bound of AD for these splays. **Preferred displacement for event Y(?); applies to independent rupture of fault splay L only. †† For the rupture of splay L only and based on mapping by Olig and Zachariasen (2008).

Evidence for noncharacteristic ruptures of intrabasin faults in the Rio Grande rift Recurrence and Slip Rates

71

⎛ 84 k.y. − 11k.y. ⎞ = 24.3 k.y. ⎟. ⎜ 3 intervals ⎝ ⎠

and J did rupture separately, this would nearly double the number of events that occurred during the past ~100 k.y. It would also imply strong temporal clustering of events and even suggest that an earthquake rupture on one splay may have triggered rupture on another splay. For comparison, and because seismic hazard analyses often do not consider complex rupture behavior alternatives (e.g., Frankel et al., 2002), we also estimated the average recurrence interval between all surface-deforming events on both splays L and J of the Hubbell Spring fault system, regardless of the type of behavior. For this estimate, we simply included event Y(?), which added another interval. This yielded a preferred average recurrence interval between all surface-deforming events of 15 k.y., with a range of 12 k.y. to 18 k.y. These recurrence interval estimates are not significantly shorter than those between simultaneous rupture events, especially given the uncertainties. Next, we used our previously calculated cumulative displacements per event (see “Rupture History and Behavior” section) to estimate cumulative slip rates for splays J and L of the Hubbell Spring fault system. Obviously, these are minimums for the entire Hubbell Spring fault system, because slip rates for other splays of the Hubbell Spring fault system are not included. However, they are still useful, particularly when looking at variations of rates through time. First, we estimated the average vertical slip rate over the past four complete seismic cycles for splays L and J. Again, assuming the oldest event (event V) occurred 70 ka, we estimated a preferred average rate of 0.18 mm/yr for splays J and L ⎛ 4.7 m + 1.7 m + 0.4 m + 3.7 m ⎞ ⎜ ⎟. 70 k.y. − 12 k.y. ⎝ ⎠

For individual recurrence intervals, we cannot improve on the original estimates of 17 and 27 k.y. by Personius and Mahan (2003) for the two youngest intervals (Fig. 7). For the interval between the first (= event V) and second (= event W) events, we estimate a preferred recurrence interval of 14 k.y., assuming event V occurred at 70 ka and event W occurred ca. 56 ka (Fig. 7). Although there is evidence for temporal clustering of surfacefaulting events on many faults in the Rio Grande rift (e.g., Foley et al., 1988; Machette, 1998; Gardner et al., 2008), we see no evidence for clustering of rupture events along splays L and J of the Hubbell Spring fault system, assuming that splay J did indeed rupture simultaneously with splay L. Individual recurrence interval estimates of 14–27 k.y. are similar to average intervals of 19 (+5/−4) k.y., and soils developed between events are consistent with these intervals. We point out, however, that although we believe that the luminescence ages, stratigraphic and pedologic correlations between trench sites, and displacement data all favor simultaneous rupture of fault splays J and L of the Hubbell Spring fault system during events Z (or fourth), X (or third), W (or second), and V (or first), we cannot preclude the possibility of independent rupture of the two splays given the large uncertainties for event ages. If splays L

Of the 10.5 m of vertical slip associated with these complete seismic cycles, over 96% (10.1 m) occurred as simultaneous rupture of splays J and L, yielding an average slip rate of 0.21 mm/yr for the past three complete seismic cycles of simultaneous rupture. In comparison, Olig and Zachariasen (2010) estimated average late Quaternary vertical slip rates of 0.2–1.0 mm/yr for the Hubbell Spring fault system based on topographic profiles across the entire system. In particular, for the topographic profile along Monterey Boulevard, just south of the Carrizo Spring site, they measured vertical displacements across splays L and J of 7.2 m and 4.0–16.0 m, respectively, with large uncertainties for splay J due to back tilting in the footwall. Next, we calculated slip rates for individual seismic cycles to see if rates have varied through time. These rates for the past four complete seismic cycles are shown in the slip rate diagram in Figure 8. Note that we consider a complete seismic cycle to include the time interval of strain buildup preceding an event and the strain release (or displacement) for that event. This is consistent with Reid’s model of elastic rebound theory for earthquake occurrence on faults. Note that since we do not know the time interval of strain buildup associated with the 2.8 m of displacement in the oldest event (0.8 m on splay L and 2 m on splay J), this is not a complete seismic cycle, and we did

Since we believe that fault splay L has ruptured both simultaneously with, and independently from, splay J of the Hubbell Spring fault system, ideally we should calculate rates of activity for both types of behavior. However, given only one observation for an independent event on splay L of the Hubbell Spring fault system, we cannot calculate recurrence intervals for this type of behavior, particularly since there are many splays of the Hubbell Spring fault system for which the paleoseismic behavior is unknown, and one of these may have actually ruptured with splay L during event Y(?). We can calculate recurrence intervals between simultaneous rupture events of splays J and L. Since timing constraints are best from the Hubbell Spring site for the youngest three events and from the Carrizo Spring site for the oldest event, we used this combination of data to determine the following intervals. Assuming a preferred age of 70 ka for event V at the Carrizo Spring site, we calculated a preferred average recurrence interval between the past four simultaneous ruptures of ~19 k.y. ⎛ 70 k.y. − 12 k.y. ⎞ = 19.3 k.y. ⎟, ⎜ 3 intervals ⎝ ⎠ with a range of 15 k.y. ⎛ 59 k.y. − 13 k.y. ⎞ = 15.3 k.y. ⎟ ⎜ ⎝ 3 intervals ⎠ to 24 k.y.

Olig et al.

not calculate an associated slip rate. Thus, the first (and oldest) complete seismic cycle is the 14 k.y. interval of strain buildup associated with 4.7 m of strain release (Fig. 8), the combined displacements for event W on fault splay L and the second event on fault splay J. The preferred slip rate for this seismic cycle is 0.34 mm/yr ⎛ ⎞ 4.7 m ⎜ ⎟. ⎝ 70 k.y. − 56 k.y. ⎠ This is nearly double the average rate, but it is still not as high as the preferred rate calculated for the youngest (fourth) complete seismic cycle, which is 0.46 mm/yr ⎛ ⎞ 3.7 m ⎜ ⎟. 20 k.y. − 12 k.y. ⎝ ⎠ In contrast, preferred rates for the second and third complete seismic cycles are as much as an order of magnitude lower, at 0.063 ⎛ ⎞ 1.7 m ⎜ ⎟ 56 k.y. − 29 k.y. ⎝ ⎠ and 0.044 ⎛ ⎞ 0.4 m ⎜ ⎟ 29 k.y. − 20 k.y. ⎝ ⎠

mm/yr, respectively (Fig. 8). We note that low slip rates are associated with both types of rupture behavior (independent and simultaneous) and result regardless of the timing uncertainties for event Y(?). Thus, although recurrence intervals have remained relatively consistent through time, cumulative slip rates for the Hubbell Spring fault appear to have varied significantly due to large variations in displacements per event, or noncharacteristic behavior. This may have important implications for hazard evaluations elsewhere in the rift. Interestingly, order-of-magnitude variations in slip rates through time have been observed on many faults throughout the Rio Grande rift (McCalpin, 1995; Machette, 1998). McCalpin (1995) observed that short-term slip rates are generally much higher than long-term rates for dozens of faults. These variations have been possibly attributed to variations in the frequency of occurrence of earthquake events (i.e., temporal clustering) or even variation in the quality or type of data being used (e.g., Wong and Olig, 1998; Machette, 1998). Results from this study indicate that noncharacteristic fault behavior can also cause order-of-magnitude variations in slip rates through time, even though recurrence intervals have remained relatively consistent. Perhaps other faults in the rift also show large slip-rate variations through time due to noncharacteristic behavior, particularly intrabasin faults with potentially complex rupture patterns like the Hubbell Spring fault system.

3.7 m

0.4 m 1.7 m

THROW (m)

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4.7 m

Date (ka) Figure 8. Plot showing the variation through time of vertical slip rates for fault splays J and L of the Hubbell Spring fault system. The first complete seismic cycle includes the time interval of strain buildup between events V (= first event) and W (= second event) and the strain (or throw) subsequently released during event W. Similarly, the second seismic cycle includes the time interval between events W and X and the cumulative throw for event X, and so on.

CS2

CS1

Locality

0–5

5–15

15–30

30–50

50–64

64–85

85–110

110–127

127–137 137–143 143–165

165–195 195–237 237–268

268–295 295–323 323–370

370–382

382–402

402–420 420–430 0–5

5–16

Bw

Ck

1Bkb

2Bkb

2Ck

3Btkb

3Bkb2

3Bkb3 4Btk 5Btk2

6K 6K2 6K3

6Btkmb 6Bkmb 7kmb2

8Btkmott

8Btkmott2

Bk Bk2 A/C

Bk

Depth (cm)

A/C

Horizon

7.5YR 7/4

7.5YR 6/5 10YR 6/4 7.5YR 5/4

7.5YR 5/6

7.5YR 5/6

7.5YR 5/4 7.5YR 5/4 7.5YR 6/4

7.5YR 8/3 7.5YR 7/3 7.5YR 6/4

7.5YR 5/4 7.5YR 7/4 7.5YR 7/4

7.5YR 5/4

7.5YR 5/4

7.5YR 4/4

7.5YR 5/4

7.5YR 5/4

7.5YR 5/4

7.5YR 5/4

Moist 7.5YR 4/3

10YR 6/3

7.5YR 7/5 7.5YR 7/6 7.5YR 6/4

7.5YR 6/5

7.5YR 6/5

7.5YR 6/4 7.5YR 7/3 7.5YR 7/3

7.5YR 8/1 7.5YR 8/2 7.5YR 7/2

7.5YR 6/4 7.5YR 7/3 7.5YR 5/3

7.5YR 6/4

7.5YR 6/4

7.5YR 5.5/4

7.5YR 5/4

7.5YR 6/4

7.5YR 5/4

10YR 5/4

25