SS suggestions 110/20/18

Spatial-temporal complexity of intro -continental intraplate seismicity: insights from geodynamic modeling and implications for seismic hazard estimation

Qingsong Li, Lunar and Planetary Institute, USRA, Houston, TX 77058

Mian Liu, Dept. of Geological Sciences, University of Missouri, Columbia, MO 65211

Seth Stein, Dept. of Earth and Planetary Sciences, Northwestern University, Evanston, IL 60208

Abstract:

Intro-c Continental intraplate seismicity has occurs in complex spatiotemporal patterns. The earthquakes are often episodic, clustered, and migrating. Hence we observe: both spatial clustering in seismic zones and scattering across large plate interiors; temporal clustering followed by long periods of quiescence; and migration of seismicity from one seismic zone (or region) to another. W To seek insight into these complexitiers, we have explored the spatiotemporal evolution of intraplate seismicity in using a 3D visco-elasto-plastic finite element model. The model simulates tectonic loading, crustal failure in earthquakes, and coseismic and postseismic stress evolution. With pre-specified perturbations of crustal strength, the model predicts various spatiotemporal patterns of seismicity at different timescales: spatial clustering (in narrow belts) and scattering (across large regions) in over hundreds of years, connected seismic belts in over thousands of years, and seismicity scattering over the whole model region over tens of thousands of years. The orientation of seismic belts coincides with the optimal failure directions determin predicted by the prescribed tectonic loading. The model results also show temporal clustering of earthquakes in seismic zones. When weak zones are included in the model, the predicted seismicity initiates within the weak zones but then extends far beyond them. If the a fault zone is weakened following a large earthquake, repeated large earthquakes can occur on the same fault zone is possible even at the absence of strong tectonic loading. These results provide useful insights into the complex spatial-temporal patterns of intraplate seismicity, and call for caution when assessing intraplate earthquake hazard. Assessment of earthquake hazard based on the limited historic record may be biased toward overestimating the risks in regions of recent large earthquakes and underestimating the risks where seismicity has been quiescent.

Introduction

The distribution of earthquakes in space and time within continental interiors is far more complex than on plate boundaries. Earthquakes on plate boundaries result from strain produced by steady relative plate motions. Hence plate boundaries remain the loci of large quasi-periodic earthquakes for long intervals until the plate boundary geometry changes. As a result, the long-term pattern of seismicity inferred from geological fault studies is consistent with that observed in large earthquakes e.g., [Cisternas, et al., 2005; Marco, et al., 1996; Weldon, et al., 2004]. Moreover, the direction and rate of geodetically observed strain accumulation is generally consistent with the mechanisms and recurrence of large earthquakes [Stein and Freymueller, 2002].

The situation is quite different within continental plate interiors, where earthquakes are clustered, episodic, and migrate (Figure 1). Paleoseismic data show that intraplate earthquakes often occur in temporal clusters on faults that remain active for some time, and then have long quiescent periods during which seismicity migrates to other faults [Crone, et al., 2003]. Thus, faults without historically recorded earthquakes, such as the Meers fault in Oklahoma [Crone and Luza, 1990], are found appear to have had prehistoric large events [Clark and McCue, 2003].

The earthquakes in continental intraplate regions may also occur in temporal clusters. For example, the earthquakes in North China’s the Weihe-Shanxi graben are temporally clustered before 1700 (Figure 2a). Similarly, The large earthquakes (M>6) in the North China Plain are clustered between 1900 and 2000 (Figure 2b). Moreover, the locus of seismicity seems migrate between faults and regions, such as from the Weihe-Shanxi graben eastward to the north China Plain in the past 200 years (Figure 1b).

The limited observations of continental intraplate seismicity show complex spatial patterns: spatial clustering of earthquakes in seismic zones and scattering in large regions (Figure 1). For example, historic and instrumentally recorded North American intracontinental earthquakes in the Central and Eastern U.S. (CEUS) appear to be concentrated in several a number of seismic zones, notably the New Madrid,and Charleston, and St. Lawrence seismic zones (Figure 1a). In addition, scattered On the other hand, seismicity occurs outside these zonesin CEUS shows certain degree of scattering (Figure 1a). Some seismic zones, such as the New Madrid seismic zone, occur along the failed rift zones, but other seismic zones do not [Schulte and Mooney, 2005]. Conversely, some prominent rifts, such as the Mid-continent rift in central North America, have little seismicity. As a result, the relation between continental intraplate seismicity and geologic structure remains uncertain and the subject of active investigation, as summarized by papers in Stein and Mazzotti [2007].

MAYBE IN FIG 1a USE THE VERSION WITH RIFTS MAPPED

The complex variation of seismicity in space and time poses a major challenge for seismic hazard estimation. It seems likely that inferring seismic hazards from historical seismicity can overestimate the hazard where n recent earthquakes have occurred and underestimate hazard elsewhere [e.g., Swafford and Stein, 2007; van Lanen and Mooney, 2007]. To reduce this difficulty, researchers are developing hazard maps that also reflect geologic structure, yielding maps with more diffuse hazards [Halchuk and Adams, 1999; Tόth, et al., 2006].

Numerous models have attempted to explain the complexities of intraplate seismicity. Generally, these models involve temporal variations of the physical properties of faults and the stress acting on them. Crone et al. [2003] suggest that temporal clustering, in which faults "turn on" to generate a series of large earthquakes and then "turn off” for along time, may reflect the evolution of pore fluid pressure in the fault zone [Sibson, 1992]. In this model, low-permeability seals form around the fault zone as stress accumulates, raising the pore pressure until an earthquake happens. Faults that weakens during earthquakes and heals during interseismic phases could also contribute the complexity of seismicity [Lyakhovsky, et al., 2001]. The migration of seismicity between faults may result from stress transfer, in which stress changes due to earthquakes affect the location of future events [Chery, et al., 2001; Mueller, et al., 2004; Rundle, et al., 2006; Stein, 1999].

Finite element model

We explored the spatiotemporal complexity of intraplate earthquakes with a 3D visco-elasto-plastic finite element model [Li and Liu, 2006; Li and Liu, 2007]. The model simulates tectonic loading, crustal failuring, and the associated coseismic and postseismic stress evolution. The t Tectonic loading is realiz simulated by applying velocity boundary conditions at the model boundaries. The c Crustal failure in an earthquake is simulated by abruptly decreasing crustal strength in an element by for a given amount when ever an element’s stress in the element reaches the failure criterion. This process releases stress at the element and causes deformation in the surrounding model domain., approximating a crustal failure event (earthquake). The p Postseismic stress evolution following the failure event is simulated with by viscous relaxation of the lower crust and the upper mantle. The coseismic and /postseismic stress changes in the failed elements that failed may increase (or decrease) stress in the neighboring regions, thus promoting (or depress inhibiting) earthquake occurrence there.

The model dimension is 2000 km by×2000 km and×60 km thick (Figure 3). A 20km thick n elasto-plastic upper layer (20km thick) simulates the upper crust and a 40km thick visco-elastic lower layer (40km thick) simulates the lower crust and the upper mantle. The eastern and western boundaries are compressed uniformly at 1 mm/yr, while the northern and southern boundaries are extended at the same rate. The bottom boundary is free-slip. Young’s moduluse is 8.75×1010 Pa and Poison’s ratio is 0.25 for both the crust and the mantle. The viscosity for the lower layer is 1×1020 Pa s. The strength of the upper layer is described by a cohesive strength of 50 MPa cohesion and a coefficient of internal friction of value of 0.4 for the internal frictional coefficient.

Figure 4 shows an example of how the predicted stress and earthquakes evolve.in an example. Figure 4a shows the stress at the beginning of a 1000-year time period, while and Figure 4b shows the stress at the end of the period. The black dots in Figure 4a show the earthquakes that occurred during the period. The earthquakes occur within regions of high stresses. After the earthquakes, stress in these regions drops below the crustal strength. The earthquakes also cause stress changes in the surrounding regions, which. However these stress changes are much smaller than the stress drops during the earthquakes.

Model results

To illustrate the impact effect of various factors on the predicted spatiotemporal patterns of intraplate seismicity, we start with a simple model with horizontally homogeneous crust and mantle, and then incrementally add more additional factors in the model.

CasCase eo 1: timescale- dependent spatiotemporal patterns of seismicity

In this case scenario, an initial a randomly pre-specified perturbation of the crustal strength is applied to an otherwise laterally homogeneous crust and mantle (Figure 3). The resulting seismicity shows both spatial clustering (in spots and belts) and scattering (in large regions) over a period of 300 a few hundred years (Figure 5a). Over a longer period, 3000 years, the earthquake clusters connect with each other to form a network of seismic zones ity (Figure 5b). In an even longer time, 30,000 years, seismicity covers the entire model domain (Figure 5c). This result is expected because the model is horizontally homogeneous. Hence seismicity is uniform over the long term, but is concentrated in seismic zones during shorter intervals.

The model results also show temporal clustering of earthquakes in each seismic zones. Figure 6 shows the predicted earthquake sequences for four 100 km2 selected regions; each is 100km×100km in size. Temporal clustering of earthquakes occurs are found in region B between 10,000 and 20,000 model years, and in region C around 5000 and 22,000 model years. The temporal clustering is caused by coseismic and postseismic stress triggering. Without coseismic and postseismic stress interaction between regions, we would expect to see periodical earthquake occurrence given that the model has constant loading and constant stress release in earthquakes.

The orientations of the predicted seismic belts, either NW-SE or NE-SW, coincide with the optimal failure directions resulting from the applied boundary conditions of: east-west compression and north-south extension. This loading condition produces a pure shear stress state in the model domain, with in which the maximum shear stress oriented is NW-SE and NE-SW oriented.

Case 2: effects of weak zones.

In the second case scenario the model does not have a prespecified perturbation of crustal strength. Instead, we explicitly incorporated in the it has three model 3 weak zones whose yield strength is 5 MPa lower than the rest of the model. Two of the zones are parallel and perpendicular to the third (Figure 7a): two are parallel to each other, and perpendicular to the third. The model does not have the pre-specified perturbation of crustal strength. The yield strength of the weak zones is 5 MPa lower than the surrounding regions. The initial stress is set to zero, and the model is constantly loaded from at the boundaries. At the beginningInitially , earthquakes occur only within the weak zones. After ~80,000 model years, stress in the surrounding regions begins to reach the strength and cause crustal failure. After another 100,000 model years, earthquake occurrence in the model enters a statistically steady state, in which the seismicity during one time period is similar to another, provided that the period is long enough (>60,000 years).

IS FIGURE 7 STARTING AT TIME ZERO OR 100,000 YEARS (STEADY STATE)?

The predicted seismicity at on different time scales is shown in Figure 7. On In a short time scale (600 years), earthquakes occur both inside and outside the weak zones (Figure 7a). The seismicity shows some spatial clustering with no clear correlation with the weak zones. Over a longer period (6000 years), the seismicity forms belts extending from the weak zone (Figure 7b). This pattern remains the same over longer periods (Figure 7c). Hence on the long time scales, seismicity is not be confined to the weak zones. Moreover, as in the previous case, the short-term seismicity pattern does not fully represent that over longer terms/

Again, we see that short-term seismicity pattern may be different from that over longer terms, and seismicity may not be confined to the weak zones.

Case 3: effects of fault weakening.

Paleoearthquake studies have found evidence of repeated ruptures on individual intraplate faults and fault zones in short periods, followed by long periods of quiescence [Crone, et al., 2003; Tuttle, et al., 2002]. The short time between the repeated earthquakes i on these intraplate faults, such as the ~500 (hundreds of years found for the recent large events in the New Madrid Zone,[Tuttle, et al., 2002]) is at odds with the low rates of deformation and far field loading in plate interiors [Calais, et al., 2005; Newman, et al., 1999]. Li et al. [2005] have shown that, a large intraplate earthquake can reduce stress in cause to the fault zone to levels below that requires for rupture remain in a stress deficiency for thousands of years. Thus, large intraplate earthquakes repeating in a short time interval would require either local loading to build up the stress, or local weakening to permit repeated failure at lowered stress levels.

A number of mechanisms for local loading and/or fault weakening have been proposedspeculated [Kenner and Segall, 2000; Pollitz, et al., 2001], although there is no direct evidence for the proposed weakening [McKenna et al., 2007]. Here w We explored the effect of fault weakening [Lyakhovsky, et al., 2001; Sibson, 1992]. In this case we explicitly incorporated one a weak fault zone (zone A in Figure 7a), whose strength is 5 MPa lower than the rest of the model domain surrounding regions. We considered here three different scenarios of fault weakening following a large fault rupture. These assume different histories of weakening but do not specified its physical mechanism.

IT’S HARD IN FIGS 8 & 9 TO SEE WHICH IS WHICH. I SUGGEST USING CONSISTENT NOTATION (A,B.C), WORDING, AND LABELS

In the first scenario, the fault zone is weakened abruptly by 0.5 MPa right immediately after the first rupture, which had a . The stress dropin the fault zone drops by of 1 Mpa.in the first event ,Stress on the fault and then gradually restores rises due to the far-field loading (Figure 8), andto reaches the new lower fault strength and causes a second fault rupture (Figure 8) in 1800 years. Because no further fault weakening occurs, iIt then takes another 3000 years for the fault zone to rupture again (Figure 9a).

In the second scenario, the strength of the fault zone weakens continuously at a rate of (1 MPa per 1000 years) after the first rupture. The stress in the weak zone recovers and is then restores and released s (Figure 8) in seismic cycles as the model predicts a series of repeated earthquakes (Figure 9b). The time interval between successive eding events is in hundreds of years.

In the third scenario, the strength of the fault zone gradually weakens in the first 2000 years following the first rupture, and then stops remains constant . The stress evolution in the first 2000 years is the same as that in the second scenario. After the weakening stops, it takes longer time for the stress to recover store (Figure 8). Thus the predicted earthquake series is identical to that in the second scenario during the first 2000 years. After that, the recurrence interval becomes increasingly longer but regular (Figure 9c).

Discussion

The strength perturbation in case -1 was randomly generated and so is hows a highly irregularity (Figure 10). The magnitude of the perturbation ranges from -0.5 to 0.5 MPa. This e purpose of the strength perturbation simulates the presence of weak zones in the ancient continental crust. is to prevent artificial effects that may be caused by uniform crustal strength. TheA model in which uniform loading was applied to a crust of uniform strength plus uniform boundary loading would predict cause a uniform failure of the whole upper crust instead of successive crustal failure at different locations and times.

Because the strength perturbation is randomly distributed, the averagaverage e strength in a large region of the model does not differ significantly from that of other regions. Consequently, the strength perturbation has only a limited effects on the distributionof seismicity. Figure 11 shows the fraction ratio of the earthquakes that occured in weak regions (negative strength perturbation) over to the total number of earthquakes in the entire model. Duringin 500-year time windows.It shows that Tthe fractions fluctuate around 0.5. In 30,000 years, 53.3% earthquakes occurred in weak regions, while 46.7% occurred in strong regions (positive strength perturbation).