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Fore-arc structure, plate coupling and isostasy in the Central Andes: Insight from gravity data modelling
Accepted Manuscript Title: Fore-arc Structure, Plate Coupling and Isostasy in the Central Andes: Insight from Gravity Data Modelling Author: Sophia Rutledge Rezene Mahatsente PII: DOI: Reference:
S0264-3707(16)30150-8 http://dx.doi.org/doi:10.1016/j.jog.2016.12.003 GEOD 1468
To appear in:
Journal of Geodynamics
Received date: Revised date: Accepted date:
3-8-2016 23-11-2016 3-12-2016
Please cite this article as: Rutledge, Sophia, Mahatsente, Rezene, Fore-arc Structure, Plate Coupling and Isostasy in the Central Andes: Insight from Gravity Data Modelling.Journal of Geodynamics http://dx.doi.org/10.1016/j.jog.2016.12.003 This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
Fore-arc Structure, Plate Coupling and Isostasy in the Central Andes: Insight from Gravity Data Modelling Sophia Rutledge and Rezene Mahatsente Dept. of Geological Sciences, The University of Alabama, Tuscaloosa, AL 35487 USA Corresponding Author: Rezene Mahatsente ([email protected])
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Abstract The central segment of the Peru-Chile subduction zone has not seen a major earthquake of similar scale to the megathrust Iquique event in 1877 (Magnitude ~ 8.8). The plate interface between the subducting and overriding plates in the central segment of the subduction zone is highly coupled and is accumulating elastic energy. Here, we assessed the locking mechanism and isostatic state of the Central Andes based on gravity models of the crust and upper mantle structure. The density models are based on satellite gravity data and are constrained by velocity models and earthquake hypocenters. The gravity models indicate a high density batholithic structure in the fore-arc, overlying the subducting Nazca plate. This high density crustal structure is pressing downward into the slab and locking the plate interface. Thus, plate coupling in the Central Andes may result from pressure exerted by high density fore-arc structures and buoyancy force on the subducting Nazca plate. The increased compressive stress closer to the trench, due to the increased contact between the subducting and overriding plates, may increase the intraplate coupling in the Central Andes. To assess the isostatic state of the Central Andes, we determined the residual topography of the region (difference between observed and isostatic topography). There is a residual topography of ~ 800 m in the western part of the Central Andes that cannot be explained by the observed crustal thicknesses. The residual topography may be attributed to mantle wedge flow and subduction of the Nazca plate. Thus, part of the observed topography in the western part of the Central Andes may be dynamically supported by mantle wedge flow below the overriding plate.
Keywords: Plate Coupling, Seismicity, Fore-arc structure and Isostasy 2
1. Introduction The Peru-Chile subduction zone, approximately 200 Myr in age, is a classic example of a mountain-forming subduction zone [Figure 1; e.g. Isacks 1988; Allmendinger et al., 1997; Oncken et al., 2006; Tassara, 2010]. The area is unique in its complexity, as the oceanic Nazca plate is subducting under the continental South American plate, evolving the margin into a highly segmented one, with along-strike variations in tectonics, subduction angle, volcanism, crustal thinning, crustal age, northward shallowing of slab and seismicity [Gutscher, et al., 2000; Lallemand et al., 2005; Tassara, 2005; Oncken et al., 2006; Moreno et al., 2008, 2009, 2011; Bilek, 2010]. The seismic history of this region involves several magnitude 7.5+ earthquakes occurring along the interface over the past 200 years, with several magnitude 8+ earthquakes occurring within the past 60 years [Bilek, 2010]. The largest earthquake to date in the region, the magnitude 9.5 Valdivia earthquake in 1960, occurred in the southern part of the subduction zone. Tsunami earthquakes, namely the 1960 and 1996 earthquakes, also occur in this region, and are characterized by slow rupture and tsunamis larger than expected for the earthquake size [Moreno et al., 2008]. The Antofagasta and Maul earthquakes, magnitudes 8.1 and 8.8, respectively, are also notable, due to the significant after-slip following the initial earthquake [Moreno, et al., 2008; Rietbrock et al., 2012; Kato and Nakagawa, 2014)]. To date, a significant section of the Peru-Chile convergent zone is building up stresses. The interseismic coupling in northern and southern Peru is significantly high indicating, elastic energy accumulation since the 1746 and 1868 earthquakes of magnitude 8.6 and 8.8 , respectively [Chlieh et al., 2011]. Similar seismic patterns have also been observed in Central Chile. The plate interface beneath Central Chile is highly coupled [Métois et al., 2013; Béjar-
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Pizarro et al., 2013]. The seismic gaps, where the plates are highly coupled, are separated by narrow zones of low coupling [Métois et al., 2013; Béjar-Pizarro et al., 2013]. The most recent seismic event in the region, the 2014 Iquique earthquake magnitude 8.2, occurred in the weakly coupled segment off the coast of Pisagua [Kato and Nakagawa, 2014]. The reasons for the seismic gaps are yet unknown, but the configuration of the slab is thought to be the main factor, based on correlations between features of the oceanic plate at the trench and along-strike changes in slab geometry and segment boundaries [Tassara et al., 2006]. Previous studies suggest northern shallowing of the slab, with different studies hypothesizing both a gradual change in slab angle and several abrupt changes in angle along strike [Gutscher, et al., 2000; Lallemand et al., 2005; Tassara, 2005; Moreno et al., 2009]. Lallemand et al. [2005] noted a correlation between strain regime and shallow dip, indicating that the areas with a low angle of dip have more compressive stress due to the increased contact area between the slab and the overriding plate. Two common hypotheses exist to describe the locking mechanisms along the slab, which have so far kept the slab from slipping to cause a major earthquake. The first is that highly buoyant oceanic features undergo subduction, pulling the slab upward into the overlying plate [Álvarez et al., 2014]. The second is that a high density body exists at the interface between the slab and the continental plate, putting downward pressure on the slab [Delouis et al., 1996 Sobiesiak et al., 2007; Tassara 2010; Schaller et al., 2015]. In this paper, we assessed the locking mechanism of plate interfaces and isostatic state of the Central Andes based on gravity data modelling and digital elevation model. The main goal is twofold: first to produce a gravity model of the Central Andes, which includes density structures of the oceanic and continental crusts, the lithosphere, and the subducted slab to a depth of 250 km, and second to use the model to analyze the effects of trench parallel crustal thickness and
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density variations on plate coupling and asperity generation (regions of high seismic moment release), as well as to understand why no major earthquake has yet occurred.. Within this paper, we present two gravity models, representative of the Central Andes subduction zone, and discuss their main features and seismic implications.
2. Geologic Setting The current configuration of the Andes Mountains is the result of several major deformation events (Figure 2). The first occurred in the Early Cretaceous period, followed by a Mid Eocene event, then, most significantly, the Late Oligocene deformation, causing the division of the Farallon plate into the Nazca and Coco plates [Tassara, 2010; Charrier et al., 2013]. This latest deformation event caused an increase in the convergence rate between the Nazca plate and the South American plate, tripling it to its current rate of 7.4 cm/yr [Tassara, 2005, 2010; Charrier et al., 2013]. Presently, the Andes cover an area of approximately 800,000 km2, with elevations reaching over 6 km [Tassara et al., 2006]. The trench, running along the continental margin, is approximately 5,900 km in length, with a maximum depth of 8 km and an average width of 64 km [Álvarez et al., 2014]. Along-strike segmentation within the subduction zone allows for the splitting of the region into four sections: the northern Andes (10°N-5°S), the Central Andes (5°S-33.5°S), the southern Andes (33.5°S-46.5°S), and the austral Andes (46.5°S-56°S). Each segment differs in topography, crustal age, volcanism, and seafloor spreading [Tassara et al., 2006]. Topographically, the region peaks in the Central Andes around 16°S, with elevations reaching over 6,000 m. North and south of this peak, the elevation descends, averaging at 3,800 m [Charrier et al., 2013 ; Tassara et al., 2006]. The crustal ages vary in an almost opposite manor,
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reaching its maximum of 48 Myr at 46.5°S and decreasing to less than 1 Myr at 20°S, then increasing again to 28 Myr at 5°S; these variations in age are largely associated with fracture zones and oceanic ridges, which cause seafloor spreading and influence slab geometry [Tassara et al., 2006: Bilek, 2010; Tassara, 2010]. The regions volcanism is segmented much like the region itself. It is divided into three major volcanic zones: the Northern Volcanic Zone (NVZ), the Central Volcanic Zone (CVZ), and the Southern Volcanic Zone (SVZ). The NVZ spans the subduction zone from approximately 8°N to 0°, and a gap between it and the CVZ is characterized by an almost flat slab. The CVZ exists from 15°S to 28°S, and the SVZ begins at 34°S and continues to the end of the subduction zone. The volcanic gap between the southern and central zones has been developing for 10 Myr and is due to the occurrence of flat sections of the subducting slab at depths between 100 and 150 km [Tassara, 2005, 2010; Tassara et al., 2006]. The Central Andes can further be divided into three segments: Altiplano (5°S to 23°S), Puna (23°S to 28°S), and the Frontal Cordillera (28°S to 34°S) [Reutter and Götze, 1994; Tassara, 2005]. The area of focus for this study is within the Altiplano segment, which varies from the coast to the inner continent. The westernmost portion is the trench, which, in this area, has very little sediment due both to long term subduction erosion of the fore-arc and to the Juan Fernandez ridge, which blocks sediment transport north of 32.5°S [Álvarez et al., 2014]. Just east of the trench is the Coastal Cordillera, which runs slightly oblique to the coastline (Figure 2). The Coastal Cordillera is characterized by smooth hills and shallow valleys with altitudes less than 1200 m, which are composed of mesozioc rock. It thins northward, disappearing around 18.5°S and reappearing at 18°S [Charrier et al., 2013]. To the east of the Coastal
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Cordillera is the Coastal Depression, a 40 to 55 km wide basin filled with volcanic sediment, with altitudes ranging from 500 to 1,000 m [Tassara, 2005, 2010; Charrier et al., 2013].
Beyond the Depression is the Fore-arc Precordillera, an area composed of Eocene magmatic and late Paleozoic felsic rocks, with elevations reaching 2,300 m [Tassara, 2005; Charrier et al., 2013]. East of the Precordillera is the Western Cordillera, where elevations reach over 6,000 m and crustal thicknesses reach over 65 km [Charrier et al., 2013]. The Altiplano flanks the Western Cordillera to the east, forming a high plateau with altitudes of 3,500 m to 4,500 m. It is comprised of a 200 km wide drainage basin in the back arc of the mountain range. The easternmost section of the Andes range is the Eastern Cordillera, which is an uplifted block of early Paleozoic sedimentary rocks bound by divergent thrust systems. It is elongated and narrow, with 4,000 m high ranges which were uplifted during the deformation of the Precordillera [Charrier et al., 2013; Tassara, 2010].
3. Data and Methods 3.1 Gravity Database Onshore and offshore gravity data, a Bougur map of which can be seen in Figure 3, is based on surface and satellite data obtained from the European Improved Gravity model of the Earth [EIGEN-6C2; Förste et al., 2012]. The EIGEN-6C2 geopotential model is a spherical harmonic representation of the Earth’s gravity field up to degree and order of 1949 (spatial resolution ca. 10 km). The model includes terrestrial and satellite gravity data from the LAGEOS (Laser Geodynamics Satellites), GRACE (Gravity Recovery And Climate Experiment) [Tapley et al, 2007; Pavlis et al, 2012] and GOCE (Gravity Field and Steady-State Ocean Circulation Explorer)
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satellite missions and is based on the WGS84 reference ellipsoid. The terrestrial data over the oceans and lands are from the altimetry-derived gravity anomalies and Earth Gravitational Model (EGM2008), respectively [Andersen and Kundsen, 1998; Pavlis et al., 2012]. There are existing terrestrial gravity data of varying quality in the Andean region. These data are included in the EGM2008 gravitational model [Pavlis et al., 2012]. We downloaded free air anomaly from the data portal of the International Centre for Global Earth Models (ICGEM: http://icgem.gfz-potsdam.de/ICGEM/), from which we then computed the complete Bouguer anomaly of the study area using Gravity Terrain Correction code [GTeC; Cella, 201]. The computation of the complete Bouguer anomaly considers spherical cap and terrain corrections up to a radius of 168 km. The corrections are based on the ETOPO - 1 global relief model of the Earth’s surface [Amante and Eakins, 2009] and the standard reduction density of 2.670 g cm-3 . The EIGEN-6C2 model is suitable for lithospheric scale modelling and can resolve geologic structures whose lateral dimensions are more than the resolution of the geopotential models. This has been demonstrated in the central Andes where a well constrained 3D density model was fitted to various geopotential models (Koether et al., 2012; Gutknecht et al., 2014).
3.2 Initial Model & Data Constraints 2.5-D gravity models were created to define the deep crust and upper mantle structure of the study area. The models run along a latitudinal line and span approximately 8 km, with a maximum depth of 250 km. Two gravity models, representative of the southern Peru and the northern Chile subduction zones are discussed here. The model in southern Peru is at 16°S and
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goes from 70°W to 80°W. The second model is in northern Chile (at 19°S) and spans from 67°W to 75°W. The locations of the models are shown in Figure 1. The densities and geometries of major structures of the initial gravity model such as the Moho, the slab and Lithosphere-Asthenosphere Boundary (LAB) are constrained by velocity models and earthquake data of several published studies in the region [Gutscher et al., 2000; Husen et al., 2000; Yuan et al., 2000; Oncken et al., 2003; Comte et al., 2004; Krabbenhöft, et al., 2004; Bilek 2010]. The thickness and density of the oceanic and continental crust were obtained from results of seismic refraction-reflection experiments and earthquake tomography (Husen et al., 2000; Oncken et al., 2003; Krabbenhöft, et al., 2004).The P-wave velocity of the oceanic crust ranges from 6.5 to 7.3 km s-1 [Oncken et al., 2003; Krabbenhöft et al., 2004]. The P-wave velocities of the continental upper and lower crust range from 5.0 to 7.0 km s-1 and from 6.7 to 7.3 km s-1, respectively [Husen et al., 2000; Oncken et al., 2003; Krabbenhöft et al., 2004]. The densities and geometries of the subducting Nazca slab and the overriding South American continental plate were constrained by results of seismic tomography and receiver function [Yuan et al., 2000; Gutscher et al., 2000; Oncken et al., 2003; Krabbenhöft et al., 2004; Comte et al., 2004]. The velocity of the oceanic and continental lithospheric mantle ranges from 7.8 to 8.15 km s-1 and from 7.6 to 8.1 km s-1, respectively [Norabuena and Snoke, 1994; Oncken et al., 2003; Krabbenhöft et al., 2004]. The velocity of the continental lithospheric mantle represents a depleted cratonic mantle beneath Central Andes [Franz and Lucasse, 2005].The density of the asthenosphere is based on P-wave travel time model [Norabuena and Snoke, 1994]. The densities of the sediments were derived from P-wave velocity of the continental margin [2.0 – 4.5 km s-1 ; Husen et al., 2000; Krabbenhöft et al., 2004]. Table 1 shows the P-wave velocities and densities of major tectonic units in the Central Andes. The density values listed in Table 1
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are derived from in situ (measured) P-wave velocity structures of Central Andes at relevant PT conditions (Pressure and Temperature) using empirical relation between P-wave velocity and density [Sobolev & Babeyko, 1994; Nafe and Drake, 1957]. The temperatures are obtained from thermal model of Central Andes [Springer, 1998], and the pressures were calculated assuming a pressure gradient of 30 MPa/km. The effect of uncertainty of the temperature and pressure values on the derived-density is negligible and does not change model results (See density variation in Table 2). The density values of the final gravity model and their tolerable variations are given in Table 2. The densities were set as contrasts relative to a background reference density model that consists of three horizontal continental layers. The first layer is 15 km thick upper crust with density value of 2.7 g cm-3, the second layer is 25 km thick and has a density value of 3.0 g cm-3, and the third layer consists of an upper mantle material down to a depth of 250 km. The third layer has a density of 3.40 g cm-3. The model is consistent with the global seismic structure of continental crust, which is based on seismic refraction profiles compiled by Christensen and Mooney [1995]. Forward modeling of the Bouguer anomaly was performed using the Grav2dc modeling software [Cooper, 1996] which makes use of Talwani’s algorithm to calculate the anomaly. We followed a 2.5-D modelling approach. The main difference between 2-D and 2.5-D is that a 2-D model has an infinite strike length, whereas a 2.5-D model has a finite strike length. We used variable strike lengths on either side of the gravity models depending on the general geological strike of the convergent zone. The modelling space, and hence the mass therein, is a small fraction of the entire Earth. This has an effect on the forward gravity modelling, because the levels of the observed and modelled
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gravity values are not identical and can result in offsets between the measured and modelled gravity values. To account for the effect of the surrounding mass on the forward gravity modelling, the dimension of the modelling space was extended to 10,000 km beyond the dimension of the 2.5-D gravity model.
4.0 Results and Discussions 4.1 Gravity Anomalies & Isostasy The Bouguer gravity anomaly map of the Central Andes indicates variations both along strike and from the ocean to the Western Cordillera (Figure 3).The gravity anomaly correlates inversely with the topography. A negative anomaly as low as -650 mGal is observed in the high cordilleras between 13°S and 14°S, where elevation is at its maximum. It increases to 0 mGal close to the shoreline, in the Coastal Cordillera, where the elevation nears and reaches sea level. The anomaly becomes positive just east of the shoreline and increases seaward, reaching 440 mGal, as the elevation decreases to below sea level and reaches depths as low as 7,670 m. The trench has an average anomaly of approximately 280 mGal, which increases northward as it gets closer to the Nazca Ridge, where the anomaly is at approximately 220 mGal. Along the Nazca ridge, the anomaly becomes positive further inward than in the area to the south of the ridge. To better understand the gravity anomalies of the subduction zone, we applied an upward continuation and wavelength filtering. First, the Bouguer gravity anomalies were upward continued to various heights. At 50 km height above the surface, most of the short wavelengths were attenuated, and the anomalies were regional in nature. In the second step, the Bouguer gravity anomalies were filtered using a low-pass filter for various cut-off values. The cut-off
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wavelength was selected based on similarity of amplitudes of gravity anomalies obtained using the upward continuation and wavelength filtering. The resulting regional gravity anomaly map (Figure 4) shows a strong negative anomaly correlating with the high elevation of the Andes Mountains, suggesting the occurrence of a deep compensating mass beneath the elevated region. The residual gravity map of the Central Andes (Figure 5) shows trench parallel positive residual gravity anomalies west of the Coastal Cordillera. These anomalies may be attributed to the subducting Nazca plate. The positive residual gravity anomalies also correlate with volcanic centers in the Fore-arc Precordillera (Figure 5). To assess the isostatic state of the Central Andes, we determined the residual topography of the region (difference between observed and isostatic topography) using digital elevation data and crustal thickness models (Figure 6). The elevation data were obtained from the ETOPO-1 global relief model [Amante & Eakins, 2009]; and the crustal thicknesses were derived from receiver function analysis and Crust1.0 model [Assumpcao et al., 2013; Laske et al., 2013]. The calculated residual topography is based on Airy isostatic model. The densities of the continental and oceanic crust for the Airy model are 2.800 g cm-3 and 2.900 g cm-3, respectively, and a density value of 3.300 g cm-3 is assumed for the mantle. The density values are based on velocity models in the region [Comte, et al., 2004; Krabbenhöft et al., 2004]. The residual topographic map of the Central Andes (Figure 6) shows positive and negative residual values, where positive and negative residuals indicate the under and over compensation conditions, respectively. The western part of the Central Andes is characterized by positive residual topography (~ 800 m) indicating that part of the Andes may be undercompensated. The crustal thickness beneath this region may not be sufficient to isostatically support the observed topography. The residual topography could be due to mantle wedge flow and subduction of the
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Nazca plate. Thus, part of the Andean region may be dynamically supported by mantle wedge flow below the overriding plate. The computation of the residual topography, which is based on the Airy isostatic principle, is a second order approximation and does not take into account the effect of the mass of the whole Earth. However, this does not change the pattern of the calculated residual topography and interpretation. The main influencing parameter is crustal thickness, and this is well constrained in this study by using crustal thickness from high resolution seismic refraction and reflection studies in the region [Comte, et al., 2004; Krabbenhöft et al., 2004].
4.2 Lithospheric Structure & Locking Mechanisms The gravity models of the subduction zone at 19°S and at 16°S based on forward modeling are shown in Figures 7a & b, respectively. They consist of the ocean, sediment, the oceanic plate, the continental plate, and the asthenosphere, and include the various bodies which make up both plates. The Nazca Plate contains a crustal body overlying a mantle lithospheric body, both of which subduct under the continental plate. The crust and the mantle bodies within the subducting Nazca plate are separated into individual segments with different densities to account for density changes associated with metamorphic reactions (Figures 7a & b). This increase reflects dehydration and subsequent densification of the subducting Nazca slab with depth .The seismic wave velocity structure of the Peru-Chile subduction zone also shows velocity increase with depth within the subducting Nazca plate (6.5- 7.3 km/s ; Krabbenhöft et al., 2004; Husen et al., 2000; Oncken et al., 2003; Lueth et al., 2004). The slab exhibits variable dip. At 19°S, the slab has a shallow dip of approximately 37° close to the trench which increases to a steeper dip of 46° below the crustal root (Figure 7b). The
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steepening slope indicates a decrease in compressive stress further from the trench. The shallow dip closer to the trench indicates more compressive stress, likely due to the increased contact area between the overriding plate and the slab. At 16°S, the slab maintains a gentle slope, dipping approximately 33° (Fig. 7a). The slightly shallower and consistent dip indicates a constant compressive stress regime, likely due to the larger width of the overlying mountain and the supporting root material. The South American continental plate is comprised of bodies representing the crust, which is separated into the upper and lower crustal sections then are further divided into right and left segments with differing densities (Figs 7a & b). The continental crust overlies the continental mantle, which is underlain by the mantle wedge and asthenosphere. Crustal thicknesses and Moho boundaries were modelled based on several published data [Yuan et al., 2000; Gutscher et al., 2000; Oncken et al., 2003; Krabbenhöft et al., 2004; Comte et al., 2004;Tassara et al el., 2006; Tassara, 2010]. The left portion of the upper continental crust in the models has a density of 2.97 g cm-3 , which is much higher than the densities of the sediment on its left and the portion of the upper crust on its right (Figures 7a & b). The high density body, a batholithic structure emplaced in the upper continental crust, and buoyancy forces may be exerting pressure on the subducting Nazca plate and thereby locking the interface between the subducting and overriding plates (Figures. 7a & b). This observation is in agreement with previous studies in the Central Andes [Delouis et al., 1996; Sobiesiak et al., 2007; Tassara, 2010; Schaller et al., 2015]. An alternative and equally convincing hypothesis for the locking mechanism in the Central Andes is that a highly buoyant oceanic features control plate coupling and asperity generation. But, this has not been tested in this study.
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The gravity models show significant crustal thickness variations in the fore-arc, further supporting the idea that dense material pressing down on the mantle locks the plates together. The crust is thinner closer to the trench and increases in depth to a maximum of 60 km as the mountain elevation above sea level reaches its maximum (Figures 7a & b). The thickening of the crust indicates that a crustal root exists below the mountain range which partially supports the weight of the mountains while pressing down onto the mantle and forcing it to lock with the subducting slab. This furthers the locking of the interface between the slab and the overriding plate. Thus, the along-strike segmentation of the overriding South American plate (trench parallel crustal thickness and density variations) may be controlling plate coupling and seismicity in Central Andes. The locations of the profiles are shown in Figure 1. The white circles are Earthquake hypocenters. Slab depths were determined based on earthquake hypocenter data from the USGS Earthquake Archive, collected through the Earthquake Hazards Program. The earthquake data consists of earthquakes events from 1950-2016. The density values of the geologic units are in g cm-3.
5. Conclusions 2.5-D gravity models of the Central Andes subduction zone are presented which, after analysis, provide a possible explanation of the densities and body geometries of the subduction zone and of the locking mechanisms that have thus far prevented a large scale earthquake within the study area. The gravity models indicate the presence of a high density fore-arc structure in the continental crust which pushes downward on the slab, causing the slab to lock with the overlying continental plate. Thus, trench parallel crustal thickness and density variations in the
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Central Andes may control plate coupling and asperity generation. Modelling of additional cross sections within the area would determine whether such high-density bodies exist elsewhere within the area. The steepening of the slab and the slab geometry shown in the gravity models further the understanding of the compressional forces within the region. The contact area between the overriding plate and the slab, as determined from gravity modelling, increases towards the trench, indicating that the fore-arc is under increased compressional stress which decreases away from the trench. The gravity anomaly appears to correlate inversely with the topography, suggesting a crustal root below the high Cordilleras that causes a low in the gravity anomaly. However, the crustal thickness in the western part of the Central Andes is not sufficient to isostatically support the observed topography. There is a positive residual topography (~ 800 m) in the western part of the Central Andes that cannot be explained by the observed crustal root. Thus, the observed topography in the western part of the Central Andes may be partially supported by mantle wedge flow below the overriding plate.
6. Acknowledgements This work has been supported by CARSCA (College Academy of Research, Scholarship, and Creative Activity, a unit of the College of Arts and Sciences, The University of Alabama) grant to R.M. We thank the three anonymous reviewers for thoughtful and constructive comments, which were of great help in the preparation of the final version of this paper.
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Förste, C., Bruinsma, S. L., Flechtner, F., Marty, J., Lemoine, J. M., Dahle, C., & Balmino, G. (2012, December). A preliminary update of the Direct approach GOCE Processing and a new release of EIGEN-6C. In AGU Fall Meeting Abstracts (Vol. 1, p. 0923). Franz, G., and Lucasse, F., (2005), 600 Ma of crustal evolution in the Central Andes: A site of crustal growth or destruction at an active continental margin?, 6th International Symposium on Andean Geodynamics (ISAG 2005, Barcelona), Extended Abstracts: 281-285 Gutknecht, B.D., Götze, H.-J., Jentzsch, G., Jahr,T., Mahatsente, R., Zeumann, S., 2014. Structure and state of stress of the Chilean subduction zone from terrestrial and satellitederived gravity and gravity gradient data. Surveys in Geophysics, DOI: 10.1007/S10712014-9296-9. Gutscher, M.-A., Spakman, W., Bijwaard, H., and Engdahl, R., 2000. Geodynamics of flat subduction: Seismicity and tomographic constraints from the Andean margin.Tectonics, 19, pp.814–833. [online] Available from: http://dx.doi.org/10.1029/1999TC001152 Husen, S., Kissling, E., Flueh, E.R., 2000. Local earthquake tomography of shallow subduction in north Chile: a combined onshore and offshore study. J. Geophys. Res. Solid Earth (19782012) 105 (B12), 28183e28198 Isacks, B., (1988), Uplift of the central Andean plateau and bending of the Bolivian orocline. J. Geophys. Res., 93, pp.3211-3231. Kato, A., Nakagawa, S., (2014). Multiple slow-slip events during a foreshock sequence of the 2014 Iquique, Chile Mw 8.1 earthquake. Geophys. Res. Lett., 41, 5420–5427, doi:10.1002/ 2014GL061138.
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Keother, N., Götze, H.-J., Gutknecht, B.D., Jahr, T., Jentzsch, G., Lücke, O.H., Mahatsente, R., Sharma, R., Zeumann, S., 2012. The seismically active Andean and Central American margins: Can satellite gravity map lithospheric structures? J. Geodyn, vol. 59, pp. 207-218. Krabbenhöff, A., J. Bialas, H. Kopp, N. Kukowsky, and C. Hu¨bscher, (2004), Crustal structure of the Peruvian continental margin from wide angle seismic studies, Geophys. J. Int., 159, 749– 764. Lallemand, S., Heuret, A. and Boutelier, D., (2005), On the relationships between slab dip, back-arc stress, upper plate absolute motion, and crustal nature in subduction zones. Geochem. Geophys. Geosyst., 6, pp1525-2027,Q09006, doi:10.1029/ 2005GC000917. [online] Available from: http://dx.doi.org/10.1029/2005GC000917 Laske, G., Masters., G., Ma, Z. and Pasyanos, M., (2013),Update on CRUST1.0 - A 1-degree Global Model of Earth's Crust, Geophys. Res. Abstracts, 15, Abstract EGU2013-2658, 2013. Métois, M., Socquet, A., Vigny,C., Carrizon, D., Peyrat, S., Delorme, A., Maureira, E., ValderasBermejo, M.-C., and Ortega, I., (2013). Revisiting the North Chile seismic gap segmentation using GPS-derived interseismic coupling. Geophys. J. Int., 194, 1283–1294, doi:10.1093/gji/ggt183. Moreno, M., Melnick, D., Rosenau, M., Bolte, J., Klotz,J., Echtler,H., Baez, J., Bataille,K., Chen,J., Bevis,M., Hase, H., Oncken, O., (2011), Heterogeneous plate locking in the South– Central Chile subduction zone: Building up the next great earthquake, Earth and Planetary Science Letters, 305(3-4), 413–424, doi:10.1016/j.epsl.2011.03.025. Moreno, M. S., Bolte, J. Klotz,J. Melnick,D., (2009), Impact of megathrust geometry on inversion of coseismic slip from geodetic data: Application to the 1960 Chile earthquake,
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Geophysical Research Letters, 36(16), doi:10.1029/2009gl039276. [online] Available from: http://dx.doi.org/10.1029/2009GL039276 Moreno, M. S., Klotz,J., Melnick,D., Echtler, H., Bataille, K., (2008), Active faulting and heterogeneous deformation across a megathrust segment boundary from GPS data, south central Chile (36-39°S), Geochemistry, Geophysics, Geosystems, 9(12), n/a–n/a, doi:10.1029/2008gc002198. [online] Available from: http://dx.doi.org/10.1029/2008GC002198 Norabuena and Snoke, (1994), Structure of the subducting Nazca Plate beneath Peru, Journal of Geophysical Research, Vol. 99, NO. B5, PAGES 9215-9226 Nafe, J. E., & Drake, C. L., (1957), Variation with depth in shallow and deep water marine sediments of porosity, density and the velocities of compressional and shear waves. Geophysics, 22(3), 523-552. Oncken, O., Chong, G., Franz, G., Giese,. P., Götze, H-J., Ramos, V., Strecker, M. and Wigger, P., (2006), The Andes – Active Subduction Orogeny. Frontiers in Earth Science Series, Vol 1. Springer-Verlag, Berlin Heidelberg New York, 567pp. Pavlis, N., Homes, S., Kenyon, S., Factor, J. (2012), The development and evaluation of the Earth Gravitational Model 2008 (EGM2008), Journal of Geophysical Research, 117(B04406), doi: , doi:10.1029/2011JB008916. Prezzi, C.B., Geotze, H.-J., Schmidt, S., 2009. 3D density model of the Central Andes. Phys. Earth Planet. Int. 177 (3e4), 217e234. Reutter K-J, Götze H-J ,(1994), Comments on the Geological and Geophysical Maps. In: Reutter, Scheuber, Wigger (eds). Tectonics of the Southern Central Andes - Structure and Evolution of an Active Continental Margin. Springer Heidelberg, pp 329-333
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Rietbrock,A., Ryder,I., Hayes,G., Haberland,C., Comte,D., Roecker,S., Lyon-Caen.,H., (2012), Aftershock seismicity of the 2010 Maule Mw=8.8, Chile, earthquake: Correlation between co-seismic slip models and aftershock distribution? Geophys. Res. Lett., Vol. 39, L08310, doi:10.1029/2012GL051308. Schaller, T., Andersen,J., Götze, H.-J., Koproch,N., Schmidt,S., Sobiesiak,M, Splettstoeßer, S., ,(2015). Segmentation of the Andean margin by isostatic models and gradients. Journal of South American Earth Sciences, 59, pp.69-85. Sobiesiak M., Meyer, U., Schmidt, S., Götze, H-J. and Krawczyk, C. M., (2007), Asperity generating upper crustal sources revealed by b value and isostatic residual anomaly grids in the area of Antofagasta, Chile. J. Geophys. Res., 112, B12308, doi:10.1029/2006JB004796. Sobolev, S.V. & Babeyko, A. Yu., 1994. Modeling of mineralogical composition, density, and elastic wave velocities in anhydrous magmatic rocks, Survey Geophys. 15, 515–544. Tapley, B., Ries, J., Bettadpur, S., Chambers, D., Cheng, M., Condi, F. and Poole, S., (2007), The GGM03 Mean Earth Gravity Model from GRACE; Eos Trans. AGU 88(52), Fall Meet. Suppl., Abstract G42A-03. Springer, M., (1999), Interpretation of heat-flow density in the Central Andes, Tectonophysics 306, 377–395 Tassara, A. (2005), Interaction between the Nazca and South American plates and formation of the Altiplano–Puna plateau: Review of a flexural analysis along the Andean margin (15°– 34°S), Tectonophysics, 399(1-4), 39–57, doi:10.1016/j.tecto.2004.12.014. [online] Available from: http://dx.doi.org/10.1016/j.tecto.2004.12.014
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Tassara, A. (2010), Control of fore-arc density structure on megathrust shear strength along the Chilean subduction zone, Tectonophysics, 495(1-2), 34–47, doi:10.1016/j.tecto.2010.06.004. [online] Available from: http://dx.doi.org/10.1016/j.tecto.2010.06.004 Tassara, A., Götze, H-J., Schmidt, S. and Hackney, R., (2006), Three-dimensional density model of the Nazca plate and the Andean continental margin. J. Geophys. Res. 111, Nr. B9, B09404, doi:10.1029/2005JB003976. [online] Available from: http://dx.doi.org/10.1029/2005JB003976. Yuan X., et al. (2000). Subduction and collision processes in the Central Andes constrained by converted seismic phases, Nature, 408, 958–961.
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Figure 1. Geotectonic setting of the Central Andes subduction zone on a digital elevation map The elevation data are from the National Oceanic and Atmospheric Administration (NOAA; http://www.ngdc.noaa.gov; Amante and Eakins, 2009). Black arrows indicate the direction of plate motion. Red triangles are for volcanos. The yellow circles indicate earthquake epicenters [USGS Earthquake Archive; https://earthquake.usgs.gov]. The black dashed lines show the locations of the 2.5-D gravity models.
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Figure 2. Simplified geologic map of the Central Andes showing major morphostructural units discussed in the text. Abbreviations: AT Altiplano, Pn Puna, FC Frontal Cordillera, CD Coastal Depression, CC Coastal Cordillera, FP Fore-arc Precordillera, WC Western Cordillera, EC Eastern Cordillera [after Charrier et al., 2013].
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Figure 3. Complete Bouguer anomaly map of the Central Andes from combined satellite and terrestrial gravity data. The gravity data are from the European Improved Gravity model of the Earth [EIGEN-6C2; Förste et al., 2012].
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Figure 4. Regional gravity anomaly map of the Central Andes.
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Figure 5. Residual gravity anomaly map of the Central Andes. The red triangles are for volcanoes.
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Figure 6. Residual topography of the Central Andes based on Airy isostasy model.
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(A)
(B)
Figure 7. 2.5-D gravity models of the Central Andes: (a) Profile at 16°S, (b) Profile at 19°S. 30
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Table 1. P-wave velocities and densities of major tectonic units in the Central Andes. The densities are derived from P-wave velocities using empirical relations [ Sobolev & Babeyko, 1994; Nafe and Drake, 1957].
Tectonic units
P- wave velocity (km s-1 )
Density (g cm-3 )
References
Continental domain: Upper & middle crust
5.0 - 7.0
2.54 - 2.97
Husen et al., 2000; Oncken et al., 2003; Krabbenhöft et al., 2004
Lower crust
6.7 -7.3
2.88 - 3.06
Husen et al., 2000; Oncken et al., 2003
Continental Mantle Lithosphere
7.6 - 8.1
3.2 - 3.33
Oncken et al., 2003
Oceanic domain: Sediment
2.0 – 4.5
1.91 – 2.5
Crust
6.5- 7.3
2.83 - 3.20
Husen et al., 200; Krabbenhöft et al., 2004 Krabbenhöft et al., 2004
Oceanic Mantle Lithosphere
7.8 - 8.15
3.22 – 3.35
Norabuena and Snoke, 1994; Krabbenhöft et al., 2004
Asthenosphere
8.15
3.35
Norabuena and Snoke, 1994
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Table 2. Density values of the final gravity model.
Tectonic units
Density (g cm-3 )
Variations (± g cm-3 )
Continental domain: Upper & middle crust Lower crust Continental Mantle Lithosphere
2.74 - 2.97 3.1 3.20 - 3.25
0.020 0.031 0.038
Oceanic domain: Sediment
2.55
0.106
Oceanic crust
2.90 - 3.20
0.025
Oceanic Mantle Lithosphere
3.20 - 3.36
0.034
Asthenosphere
3.35
0.023
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S0264-3707(16)30150-8 http://dx.doi.org/doi:10.1016/j.jog.2016.12.003 GEOD 1468
To appear in:
Journal of Geodynamics
Received date: Revised date: Accepted date:
3-8-2016 23-11-2016 3-12-2016
Please cite this article as: Rutledge, Sophia, Mahatsente, Rezene, Fore-arc Structure, Plate Coupling and Isostasy in the Central Andes: Insight from Gravity Data Modelling.Journal of Geodynamics http://dx.doi.org/10.1016/j.jog.2016.12.003 This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
Fore-arc Structure, Plate Coupling and Isostasy in the Central Andes: Insight from Gravity Data Modelling Sophia Rutledge and Rezene Mahatsente Dept. of Geological Sciences, The University of Alabama, Tuscaloosa, AL 35487 USA Corresponding Author: Rezene Mahatsente ([email protected])
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Abstract The central segment of the Peru-Chile subduction zone has not seen a major earthquake of similar scale to the megathrust Iquique event in 1877 (Magnitude ~ 8.8). The plate interface between the subducting and overriding plates in the central segment of the subduction zone is highly coupled and is accumulating elastic energy. Here, we assessed the locking mechanism and isostatic state of the Central Andes based on gravity models of the crust and upper mantle structure. The density models are based on satellite gravity data and are constrained by velocity models and earthquake hypocenters. The gravity models indicate a high density batholithic structure in the fore-arc, overlying the subducting Nazca plate. This high density crustal structure is pressing downward into the slab and locking the plate interface. Thus, plate coupling in the Central Andes may result from pressure exerted by high density fore-arc structures and buoyancy force on the subducting Nazca plate. The increased compressive stress closer to the trench, due to the increased contact between the subducting and overriding plates, may increase the intraplate coupling in the Central Andes. To assess the isostatic state of the Central Andes, we determined the residual topography of the region (difference between observed and isostatic topography). There is a residual topography of ~ 800 m in the western part of the Central Andes that cannot be explained by the observed crustal thicknesses. The residual topography may be attributed to mantle wedge flow and subduction of the Nazca plate. Thus, part of the observed topography in the western part of the Central Andes may be dynamically supported by mantle wedge flow below the overriding plate.
Keywords: Plate Coupling, Seismicity, Fore-arc structure and Isostasy 2
1. Introduction The Peru-Chile subduction zone, approximately 200 Myr in age, is a classic example of a mountain-forming subduction zone [Figure 1; e.g. Isacks 1988; Allmendinger et al., 1997; Oncken et al., 2006; Tassara, 2010]. The area is unique in its complexity, as the oceanic Nazca plate is subducting under the continental South American plate, evolving the margin into a highly segmented one, with along-strike variations in tectonics, subduction angle, volcanism, crustal thinning, crustal age, northward shallowing of slab and seismicity [Gutscher, et al., 2000; Lallemand et al., 2005; Tassara, 2005; Oncken et al., 2006; Moreno et al., 2008, 2009, 2011; Bilek, 2010]. The seismic history of this region involves several magnitude 7.5+ earthquakes occurring along the interface over the past 200 years, with several magnitude 8+ earthquakes occurring within the past 60 years [Bilek, 2010]. The largest earthquake to date in the region, the magnitude 9.5 Valdivia earthquake in 1960, occurred in the southern part of the subduction zone. Tsunami earthquakes, namely the 1960 and 1996 earthquakes, also occur in this region, and are characterized by slow rupture and tsunamis larger than expected for the earthquake size [Moreno et al., 2008]. The Antofagasta and Maul earthquakes, magnitudes 8.1 and 8.8, respectively, are also notable, due to the significant after-slip following the initial earthquake [Moreno, et al., 2008; Rietbrock et al., 2012; Kato and Nakagawa, 2014)]. To date, a significant section of the Peru-Chile convergent zone is building up stresses. The interseismic coupling in northern and southern Peru is significantly high indicating, elastic energy accumulation since the 1746 and 1868 earthquakes of magnitude 8.6 and 8.8 , respectively [Chlieh et al., 2011]. Similar seismic patterns have also been observed in Central Chile. The plate interface beneath Central Chile is highly coupled [Métois et al., 2013; Béjar-
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Pizarro et al., 2013]. The seismic gaps, where the plates are highly coupled, are separated by narrow zones of low coupling [Métois et al., 2013; Béjar-Pizarro et al., 2013]. The most recent seismic event in the region, the 2014 Iquique earthquake magnitude 8.2, occurred in the weakly coupled segment off the coast of Pisagua [Kato and Nakagawa, 2014]. The reasons for the seismic gaps are yet unknown, but the configuration of the slab is thought to be the main factor, based on correlations between features of the oceanic plate at the trench and along-strike changes in slab geometry and segment boundaries [Tassara et al., 2006]. Previous studies suggest northern shallowing of the slab, with different studies hypothesizing both a gradual change in slab angle and several abrupt changes in angle along strike [Gutscher, et al., 2000; Lallemand et al., 2005; Tassara, 2005; Moreno et al., 2009]. Lallemand et al. [2005] noted a correlation between strain regime and shallow dip, indicating that the areas with a low angle of dip have more compressive stress due to the increased contact area between the slab and the overriding plate. Two common hypotheses exist to describe the locking mechanisms along the slab, which have so far kept the slab from slipping to cause a major earthquake. The first is that highly buoyant oceanic features undergo subduction, pulling the slab upward into the overlying plate [Álvarez et al., 2014]. The second is that a high density body exists at the interface between the slab and the continental plate, putting downward pressure on the slab [Delouis et al., 1996 Sobiesiak et al., 2007; Tassara 2010; Schaller et al., 2015]. In this paper, we assessed the locking mechanism of plate interfaces and isostatic state of the Central Andes based on gravity data modelling and digital elevation model. The main goal is twofold: first to produce a gravity model of the Central Andes, which includes density structures of the oceanic and continental crusts, the lithosphere, and the subducted slab to a depth of 250 km, and second to use the model to analyze the effects of trench parallel crustal thickness and
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density variations on plate coupling and asperity generation (regions of high seismic moment release), as well as to understand why no major earthquake has yet occurred.. Within this paper, we present two gravity models, representative of the Central Andes subduction zone, and discuss their main features and seismic implications.
2. Geologic Setting The current configuration of the Andes Mountains is the result of several major deformation events (Figure 2). The first occurred in the Early Cretaceous period, followed by a Mid Eocene event, then, most significantly, the Late Oligocene deformation, causing the division of the Farallon plate into the Nazca and Coco plates [Tassara, 2010; Charrier et al., 2013]. This latest deformation event caused an increase in the convergence rate between the Nazca plate and the South American plate, tripling it to its current rate of 7.4 cm/yr [Tassara, 2005, 2010; Charrier et al., 2013]. Presently, the Andes cover an area of approximately 800,000 km2, with elevations reaching over 6 km [Tassara et al., 2006]. The trench, running along the continental margin, is approximately 5,900 km in length, with a maximum depth of 8 km and an average width of 64 km [Álvarez et al., 2014]. Along-strike segmentation within the subduction zone allows for the splitting of the region into four sections: the northern Andes (10°N-5°S), the Central Andes (5°S-33.5°S), the southern Andes (33.5°S-46.5°S), and the austral Andes (46.5°S-56°S). Each segment differs in topography, crustal age, volcanism, and seafloor spreading [Tassara et al., 2006]. Topographically, the region peaks in the Central Andes around 16°S, with elevations reaching over 6,000 m. North and south of this peak, the elevation descends, averaging at 3,800 m [Charrier et al., 2013 ; Tassara et al., 2006]. The crustal ages vary in an almost opposite manor,
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reaching its maximum of 48 Myr at 46.5°S and decreasing to less than 1 Myr at 20°S, then increasing again to 28 Myr at 5°S; these variations in age are largely associated with fracture zones and oceanic ridges, which cause seafloor spreading and influence slab geometry [Tassara et al., 2006: Bilek, 2010; Tassara, 2010]. The regions volcanism is segmented much like the region itself. It is divided into three major volcanic zones: the Northern Volcanic Zone (NVZ), the Central Volcanic Zone (CVZ), and the Southern Volcanic Zone (SVZ). The NVZ spans the subduction zone from approximately 8°N to 0°, and a gap between it and the CVZ is characterized by an almost flat slab. The CVZ exists from 15°S to 28°S, and the SVZ begins at 34°S and continues to the end of the subduction zone. The volcanic gap between the southern and central zones has been developing for 10 Myr and is due to the occurrence of flat sections of the subducting slab at depths between 100 and 150 km [Tassara, 2005, 2010; Tassara et al., 2006]. The Central Andes can further be divided into three segments: Altiplano (5°S to 23°S), Puna (23°S to 28°S), and the Frontal Cordillera (28°S to 34°S) [Reutter and Götze, 1994; Tassara, 2005]. The area of focus for this study is within the Altiplano segment, which varies from the coast to the inner continent. The westernmost portion is the trench, which, in this area, has very little sediment due both to long term subduction erosion of the fore-arc and to the Juan Fernandez ridge, which blocks sediment transport north of 32.5°S [Álvarez et al., 2014]. Just east of the trench is the Coastal Cordillera, which runs slightly oblique to the coastline (Figure 2). The Coastal Cordillera is characterized by smooth hills and shallow valleys with altitudes less than 1200 m, which are composed of mesozioc rock. It thins northward, disappearing around 18.5°S and reappearing at 18°S [Charrier et al., 2013]. To the east of the Coastal
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Cordillera is the Coastal Depression, a 40 to 55 km wide basin filled with volcanic sediment, with altitudes ranging from 500 to 1,000 m [Tassara, 2005, 2010; Charrier et al., 2013].
Beyond the Depression is the Fore-arc Precordillera, an area composed of Eocene magmatic and late Paleozoic felsic rocks, with elevations reaching 2,300 m [Tassara, 2005; Charrier et al., 2013]. East of the Precordillera is the Western Cordillera, where elevations reach over 6,000 m and crustal thicknesses reach over 65 km [Charrier et al., 2013]. The Altiplano flanks the Western Cordillera to the east, forming a high plateau with altitudes of 3,500 m to 4,500 m. It is comprised of a 200 km wide drainage basin in the back arc of the mountain range. The easternmost section of the Andes range is the Eastern Cordillera, which is an uplifted block of early Paleozoic sedimentary rocks bound by divergent thrust systems. It is elongated and narrow, with 4,000 m high ranges which were uplifted during the deformation of the Precordillera [Charrier et al., 2013; Tassara, 2010].
3. Data and Methods 3.1 Gravity Database Onshore and offshore gravity data, a Bougur map of which can be seen in Figure 3, is based on surface and satellite data obtained from the European Improved Gravity model of the Earth [EIGEN-6C2; Förste et al., 2012]. The EIGEN-6C2 geopotential model is a spherical harmonic representation of the Earth’s gravity field up to degree and order of 1949 (spatial resolution ca. 10 km). The model includes terrestrial and satellite gravity data from the LAGEOS (Laser Geodynamics Satellites), GRACE (Gravity Recovery And Climate Experiment) [Tapley et al, 2007; Pavlis et al, 2012] and GOCE (Gravity Field and Steady-State Ocean Circulation Explorer)
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satellite missions and is based on the WGS84 reference ellipsoid. The terrestrial data over the oceans and lands are from the altimetry-derived gravity anomalies and Earth Gravitational Model (EGM2008), respectively [Andersen and Kundsen, 1998; Pavlis et al., 2012]. There are existing terrestrial gravity data of varying quality in the Andean region. These data are included in the EGM2008 gravitational model [Pavlis et al., 2012]. We downloaded free air anomaly from the data portal of the International Centre for Global Earth Models (ICGEM: http://icgem.gfz-potsdam.de/ICGEM/), from which we then computed the complete Bouguer anomaly of the study area using Gravity Terrain Correction code [GTeC; Cella, 201]. The computation of the complete Bouguer anomaly considers spherical cap and terrain corrections up to a radius of 168 km. The corrections are based on the ETOPO - 1 global relief model of the Earth’s surface [Amante and Eakins, 2009] and the standard reduction density of 2.670 g cm-3 . The EIGEN-6C2 model is suitable for lithospheric scale modelling and can resolve geologic structures whose lateral dimensions are more than the resolution of the geopotential models. This has been demonstrated in the central Andes where a well constrained 3D density model was fitted to various geopotential models (Koether et al., 2012; Gutknecht et al., 2014).
3.2 Initial Model & Data Constraints 2.5-D gravity models were created to define the deep crust and upper mantle structure of the study area. The models run along a latitudinal line and span approximately 8 km, with a maximum depth of 250 km. Two gravity models, representative of the southern Peru and the northern Chile subduction zones are discussed here. The model in southern Peru is at 16°S and
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goes from 70°W to 80°W. The second model is in northern Chile (at 19°S) and spans from 67°W to 75°W. The locations of the models are shown in Figure 1. The densities and geometries of major structures of the initial gravity model such as the Moho, the slab and Lithosphere-Asthenosphere Boundary (LAB) are constrained by velocity models and earthquake data of several published studies in the region [Gutscher et al., 2000; Husen et al., 2000; Yuan et al., 2000; Oncken et al., 2003; Comte et al., 2004; Krabbenhöft, et al., 2004; Bilek 2010]. The thickness and density of the oceanic and continental crust were obtained from results of seismic refraction-reflection experiments and earthquake tomography (Husen et al., 2000; Oncken et al., 2003; Krabbenhöft, et al., 2004).The P-wave velocity of the oceanic crust ranges from 6.5 to 7.3 km s-1 [Oncken et al., 2003; Krabbenhöft et al., 2004]. The P-wave velocities of the continental upper and lower crust range from 5.0 to 7.0 km s-1 and from 6.7 to 7.3 km s-1, respectively [Husen et al., 2000; Oncken et al., 2003; Krabbenhöft et al., 2004]. The densities and geometries of the subducting Nazca slab and the overriding South American continental plate were constrained by results of seismic tomography and receiver function [Yuan et al., 2000; Gutscher et al., 2000; Oncken et al., 2003; Krabbenhöft et al., 2004; Comte et al., 2004]. The velocity of the oceanic and continental lithospheric mantle ranges from 7.8 to 8.15 km s-1 and from 7.6 to 8.1 km s-1, respectively [Norabuena and Snoke, 1994; Oncken et al., 2003; Krabbenhöft et al., 2004]. The velocity of the continental lithospheric mantle represents a depleted cratonic mantle beneath Central Andes [Franz and Lucasse, 2005].The density of the asthenosphere is based on P-wave travel time model [Norabuena and Snoke, 1994]. The densities of the sediments were derived from P-wave velocity of the continental margin [2.0 – 4.5 km s-1 ; Husen et al., 2000; Krabbenhöft et al., 2004]. Table 1 shows the P-wave velocities and densities of major tectonic units in the Central Andes. The density values listed in Table 1
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are derived from in situ (measured) P-wave velocity structures of Central Andes at relevant PT conditions (Pressure and Temperature) using empirical relation between P-wave velocity and density [Sobolev & Babeyko, 1994; Nafe and Drake, 1957]. The temperatures are obtained from thermal model of Central Andes [Springer, 1998], and the pressures were calculated assuming a pressure gradient of 30 MPa/km. The effect of uncertainty of the temperature and pressure values on the derived-density is negligible and does not change model results (See density variation in Table 2). The density values of the final gravity model and their tolerable variations are given in Table 2. The densities were set as contrasts relative to a background reference density model that consists of three horizontal continental layers. The first layer is 15 km thick upper crust with density value of 2.7 g cm-3, the second layer is 25 km thick and has a density value of 3.0 g cm-3, and the third layer consists of an upper mantle material down to a depth of 250 km. The third layer has a density of 3.40 g cm-3. The model is consistent with the global seismic structure of continental crust, which is based on seismic refraction profiles compiled by Christensen and Mooney [1995]. Forward modeling of the Bouguer anomaly was performed using the Grav2dc modeling software [Cooper, 1996] which makes use of Talwani’s algorithm to calculate the anomaly. We followed a 2.5-D modelling approach. The main difference between 2-D and 2.5-D is that a 2-D model has an infinite strike length, whereas a 2.5-D model has a finite strike length. We used variable strike lengths on either side of the gravity models depending on the general geological strike of the convergent zone. The modelling space, and hence the mass therein, is a small fraction of the entire Earth. This has an effect on the forward gravity modelling, because the levels of the observed and modelled
10
gravity values are not identical and can result in offsets between the measured and modelled gravity values. To account for the effect of the surrounding mass on the forward gravity modelling, the dimension of the modelling space was extended to 10,000 km beyond the dimension of the 2.5-D gravity model.
4.0 Results and Discussions 4.1 Gravity Anomalies & Isostasy The Bouguer gravity anomaly map of the Central Andes indicates variations both along strike and from the ocean to the Western Cordillera (Figure 3).The gravity anomaly correlates inversely with the topography. A negative anomaly as low as -650 mGal is observed in the high cordilleras between 13°S and 14°S, where elevation is at its maximum. It increases to 0 mGal close to the shoreline, in the Coastal Cordillera, where the elevation nears and reaches sea level. The anomaly becomes positive just east of the shoreline and increases seaward, reaching 440 mGal, as the elevation decreases to below sea level and reaches depths as low as 7,670 m. The trench has an average anomaly of approximately 280 mGal, which increases northward as it gets closer to the Nazca Ridge, where the anomaly is at approximately 220 mGal. Along the Nazca ridge, the anomaly becomes positive further inward than in the area to the south of the ridge. To better understand the gravity anomalies of the subduction zone, we applied an upward continuation and wavelength filtering. First, the Bouguer gravity anomalies were upward continued to various heights. At 50 km height above the surface, most of the short wavelengths were attenuated, and the anomalies were regional in nature. In the second step, the Bouguer gravity anomalies were filtered using a low-pass filter for various cut-off values. The cut-off
11
wavelength was selected based on similarity of amplitudes of gravity anomalies obtained using the upward continuation and wavelength filtering. The resulting regional gravity anomaly map (Figure 4) shows a strong negative anomaly correlating with the high elevation of the Andes Mountains, suggesting the occurrence of a deep compensating mass beneath the elevated region. The residual gravity map of the Central Andes (Figure 5) shows trench parallel positive residual gravity anomalies west of the Coastal Cordillera. These anomalies may be attributed to the subducting Nazca plate. The positive residual gravity anomalies also correlate with volcanic centers in the Fore-arc Precordillera (Figure 5). To assess the isostatic state of the Central Andes, we determined the residual topography of the region (difference between observed and isostatic topography) using digital elevation data and crustal thickness models (Figure 6). The elevation data were obtained from the ETOPO-1 global relief model [Amante & Eakins, 2009]; and the crustal thicknesses were derived from receiver function analysis and Crust1.0 model [Assumpcao et al., 2013; Laske et al., 2013]. The calculated residual topography is based on Airy isostatic model. The densities of the continental and oceanic crust for the Airy model are 2.800 g cm-3 and 2.900 g cm-3, respectively, and a density value of 3.300 g cm-3 is assumed for the mantle. The density values are based on velocity models in the region [Comte, et al., 2004; Krabbenhöft et al., 2004]. The residual topographic map of the Central Andes (Figure 6) shows positive and negative residual values, where positive and negative residuals indicate the under and over compensation conditions, respectively. The western part of the Central Andes is characterized by positive residual topography (~ 800 m) indicating that part of the Andes may be undercompensated. The crustal thickness beneath this region may not be sufficient to isostatically support the observed topography. The residual topography could be due to mantle wedge flow and subduction of the
12
Nazca plate. Thus, part of the Andean region may be dynamically supported by mantle wedge flow below the overriding plate. The computation of the residual topography, which is based on the Airy isostatic principle, is a second order approximation and does not take into account the effect of the mass of the whole Earth. However, this does not change the pattern of the calculated residual topography and interpretation. The main influencing parameter is crustal thickness, and this is well constrained in this study by using crustal thickness from high resolution seismic refraction and reflection studies in the region [Comte, et al., 2004; Krabbenhöft et al., 2004].
4.2 Lithospheric Structure & Locking Mechanisms The gravity models of the subduction zone at 19°S and at 16°S based on forward modeling are shown in Figures 7a & b, respectively. They consist of the ocean, sediment, the oceanic plate, the continental plate, and the asthenosphere, and include the various bodies which make up both plates. The Nazca Plate contains a crustal body overlying a mantle lithospheric body, both of which subduct under the continental plate. The crust and the mantle bodies within the subducting Nazca plate are separated into individual segments with different densities to account for density changes associated with metamorphic reactions (Figures 7a & b). This increase reflects dehydration and subsequent densification of the subducting Nazca slab with depth .The seismic wave velocity structure of the Peru-Chile subduction zone also shows velocity increase with depth within the subducting Nazca plate (6.5- 7.3 km/s ; Krabbenhöft et al., 2004; Husen et al., 2000; Oncken et al., 2003; Lueth et al., 2004). The slab exhibits variable dip. At 19°S, the slab has a shallow dip of approximately 37° close to the trench which increases to a steeper dip of 46° below the crustal root (Figure 7b). The
13
steepening slope indicates a decrease in compressive stress further from the trench. The shallow dip closer to the trench indicates more compressive stress, likely due to the increased contact area between the overriding plate and the slab. At 16°S, the slab maintains a gentle slope, dipping approximately 33° (Fig. 7a). The slightly shallower and consistent dip indicates a constant compressive stress regime, likely due to the larger width of the overlying mountain and the supporting root material. The South American continental plate is comprised of bodies representing the crust, which is separated into the upper and lower crustal sections then are further divided into right and left segments with differing densities (Figs 7a & b). The continental crust overlies the continental mantle, which is underlain by the mantle wedge and asthenosphere. Crustal thicknesses and Moho boundaries were modelled based on several published data [Yuan et al., 2000; Gutscher et al., 2000; Oncken et al., 2003; Krabbenhöft et al., 2004; Comte et al., 2004;Tassara et al el., 2006; Tassara, 2010]. The left portion of the upper continental crust in the models has a density of 2.97 g cm-3 , which is much higher than the densities of the sediment on its left and the portion of the upper crust on its right (Figures 7a & b). The high density body, a batholithic structure emplaced in the upper continental crust, and buoyancy forces may be exerting pressure on the subducting Nazca plate and thereby locking the interface between the subducting and overriding plates (Figures. 7a & b). This observation is in agreement with previous studies in the Central Andes [Delouis et al., 1996; Sobiesiak et al., 2007; Tassara, 2010; Schaller et al., 2015]. An alternative and equally convincing hypothesis for the locking mechanism in the Central Andes is that a highly buoyant oceanic features control plate coupling and asperity generation. But, this has not been tested in this study.
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The gravity models show significant crustal thickness variations in the fore-arc, further supporting the idea that dense material pressing down on the mantle locks the plates together. The crust is thinner closer to the trench and increases in depth to a maximum of 60 km as the mountain elevation above sea level reaches its maximum (Figures 7a & b). The thickening of the crust indicates that a crustal root exists below the mountain range which partially supports the weight of the mountains while pressing down onto the mantle and forcing it to lock with the subducting slab. This furthers the locking of the interface between the slab and the overriding plate. Thus, the along-strike segmentation of the overriding South American plate (trench parallel crustal thickness and density variations) may be controlling plate coupling and seismicity in Central Andes. The locations of the profiles are shown in Figure 1. The white circles are Earthquake hypocenters. Slab depths were determined based on earthquake hypocenter data from the USGS Earthquake Archive, collected through the Earthquake Hazards Program. The earthquake data consists of earthquakes events from 1950-2016. The density values of the geologic units are in g cm-3.
5. Conclusions 2.5-D gravity models of the Central Andes subduction zone are presented which, after analysis, provide a possible explanation of the densities and body geometries of the subduction zone and of the locking mechanisms that have thus far prevented a large scale earthquake within the study area. The gravity models indicate the presence of a high density fore-arc structure in the continental crust which pushes downward on the slab, causing the slab to lock with the overlying continental plate. Thus, trench parallel crustal thickness and density variations in the
15
Central Andes may control plate coupling and asperity generation. Modelling of additional cross sections within the area would determine whether such high-density bodies exist elsewhere within the area. The steepening of the slab and the slab geometry shown in the gravity models further the understanding of the compressional forces within the region. The contact area between the overriding plate and the slab, as determined from gravity modelling, increases towards the trench, indicating that the fore-arc is under increased compressional stress which decreases away from the trench. The gravity anomaly appears to correlate inversely with the topography, suggesting a crustal root below the high Cordilleras that causes a low in the gravity anomaly. However, the crustal thickness in the western part of the Central Andes is not sufficient to isostatically support the observed topography. There is a positive residual topography (~ 800 m) in the western part of the Central Andes that cannot be explained by the observed crustal root. Thus, the observed topography in the western part of the Central Andes may be partially supported by mantle wedge flow below the overriding plate.
6. Acknowledgements This work has been supported by CARSCA (College Academy of Research, Scholarship, and Creative Activity, a unit of the College of Arts and Sciences, The University of Alabama) grant to R.M. We thank the three anonymous reviewers for thoughtful and constructive comments, which were of great help in the preparation of the final version of this paper.
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Keother, N., Götze, H.-J., Gutknecht, B.D., Jahr, T., Jentzsch, G., Lücke, O.H., Mahatsente, R., Sharma, R., Zeumann, S., 2012. The seismically active Andean and Central American margins: Can satellite gravity map lithospheric structures? J. Geodyn, vol. 59, pp. 207-218. Krabbenhöff, A., J. Bialas, H. Kopp, N. Kukowsky, and C. Hu¨bscher, (2004), Crustal structure of the Peruvian continental margin from wide angle seismic studies, Geophys. J. Int., 159, 749– 764. Lallemand, S., Heuret, A. and Boutelier, D., (2005), On the relationships between slab dip, back-arc stress, upper plate absolute motion, and crustal nature in subduction zones. Geochem. Geophys. Geosyst., 6, pp1525-2027,Q09006, doi:10.1029/ 2005GC000917. [online] Available from: http://dx.doi.org/10.1029/2005GC000917 Laske, G., Masters., G., Ma, Z. and Pasyanos, M., (2013),Update on CRUST1.0 - A 1-degree Global Model of Earth's Crust, Geophys. Res. Abstracts, 15, Abstract EGU2013-2658, 2013. Métois, M., Socquet, A., Vigny,C., Carrizon, D., Peyrat, S., Delorme, A., Maureira, E., ValderasBermejo, M.-C., and Ortega, I., (2013). Revisiting the North Chile seismic gap segmentation using GPS-derived interseismic coupling. Geophys. J. Int., 194, 1283–1294, doi:10.1093/gji/ggt183. Moreno, M., Melnick, D., Rosenau, M., Bolte, J., Klotz,J., Echtler,H., Baez, J., Bataille,K., Chen,J., Bevis,M., Hase, H., Oncken, O., (2011), Heterogeneous plate locking in the South– Central Chile subduction zone: Building up the next great earthquake, Earth and Planetary Science Letters, 305(3-4), 413–424, doi:10.1016/j.epsl.2011.03.025. Moreno, M. S., Bolte, J. Klotz,J. Melnick,D., (2009), Impact of megathrust geometry on inversion of coseismic slip from geodetic data: Application to the 1960 Chile earthquake,
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Geophysical Research Letters, 36(16), doi:10.1029/2009gl039276. [online] Available from: http://dx.doi.org/10.1029/2009GL039276 Moreno, M. S., Klotz,J., Melnick,D., Echtler, H., Bataille, K., (2008), Active faulting and heterogeneous deformation across a megathrust segment boundary from GPS data, south central Chile (36-39°S), Geochemistry, Geophysics, Geosystems, 9(12), n/a–n/a, doi:10.1029/2008gc002198. [online] Available from: http://dx.doi.org/10.1029/2008GC002198 Norabuena and Snoke, (1994), Structure of the subducting Nazca Plate beneath Peru, Journal of Geophysical Research, Vol. 99, NO. B5, PAGES 9215-9226 Nafe, J. E., & Drake, C. L., (1957), Variation with depth in shallow and deep water marine sediments of porosity, density and the velocities of compressional and shear waves. Geophysics, 22(3), 523-552. Oncken, O., Chong, G., Franz, G., Giese,. P., Götze, H-J., Ramos, V., Strecker, M. and Wigger, P., (2006), The Andes – Active Subduction Orogeny. Frontiers in Earth Science Series, Vol 1. Springer-Verlag, Berlin Heidelberg New York, 567pp. Pavlis, N., Homes, S., Kenyon, S., Factor, J. (2012), The development and evaluation of the Earth Gravitational Model 2008 (EGM2008), Journal of Geophysical Research, 117(B04406), doi: , doi:10.1029/2011JB008916. Prezzi, C.B., Geotze, H.-J., Schmidt, S., 2009. 3D density model of the Central Andes. Phys. Earth Planet. Int. 177 (3e4), 217e234. Reutter K-J, Götze H-J ,(1994), Comments on the Geological and Geophysical Maps. In: Reutter, Scheuber, Wigger (eds). Tectonics of the Southern Central Andes - Structure and Evolution of an Active Continental Margin. Springer Heidelberg, pp 329-333
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Rietbrock,A., Ryder,I., Hayes,G., Haberland,C., Comte,D., Roecker,S., Lyon-Caen.,H., (2012), Aftershock seismicity of the 2010 Maule Mw=8.8, Chile, earthquake: Correlation between co-seismic slip models and aftershock distribution? Geophys. Res. Lett., Vol. 39, L08310, doi:10.1029/2012GL051308. Schaller, T., Andersen,J., Götze, H.-J., Koproch,N., Schmidt,S., Sobiesiak,M, Splettstoeßer, S., ,(2015). Segmentation of the Andean margin by isostatic models and gradients. Journal of South American Earth Sciences, 59, pp.69-85. Sobiesiak M., Meyer, U., Schmidt, S., Götze, H-J. and Krawczyk, C. M., (2007), Asperity generating upper crustal sources revealed by b value and isostatic residual anomaly grids in the area of Antofagasta, Chile. J. Geophys. Res., 112, B12308, doi:10.1029/2006JB004796. Sobolev, S.V. & Babeyko, A. Yu., 1994. Modeling of mineralogical composition, density, and elastic wave velocities in anhydrous magmatic rocks, Survey Geophys. 15, 515–544. Tapley, B., Ries, J., Bettadpur, S., Chambers, D., Cheng, M., Condi, F. and Poole, S., (2007), The GGM03 Mean Earth Gravity Model from GRACE; Eos Trans. AGU 88(52), Fall Meet. Suppl., Abstract G42A-03. Springer, M., (1999), Interpretation of heat-flow density in the Central Andes, Tectonophysics 306, 377–395 Tassara, A. (2005), Interaction between the Nazca and South American plates and formation of the Altiplano–Puna plateau: Review of a flexural analysis along the Andean margin (15°– 34°S), Tectonophysics, 399(1-4), 39–57, doi:10.1016/j.tecto.2004.12.014. [online] Available from: http://dx.doi.org/10.1016/j.tecto.2004.12.014
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Tassara, A. (2010), Control of fore-arc density structure on megathrust shear strength along the Chilean subduction zone, Tectonophysics, 495(1-2), 34–47, doi:10.1016/j.tecto.2010.06.004. [online] Available from: http://dx.doi.org/10.1016/j.tecto.2010.06.004 Tassara, A., Götze, H-J., Schmidt, S. and Hackney, R., (2006), Three-dimensional density model of the Nazca plate and the Andean continental margin. J. Geophys. Res. 111, Nr. B9, B09404, doi:10.1029/2005JB003976. [online] Available from: http://dx.doi.org/10.1029/2005JB003976. Yuan X., et al. (2000). Subduction and collision processes in the Central Andes constrained by converted seismic phases, Nature, 408, 958–961.
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Figure 1. Geotectonic setting of the Central Andes subduction zone on a digital elevation map The elevation data are from the National Oceanic and Atmospheric Administration (NOAA; http://www.ngdc.noaa.gov; Amante and Eakins, 2009). Black arrows indicate the direction of plate motion. Red triangles are for volcanos. The yellow circles indicate earthquake epicenters [USGS Earthquake Archive; https://earthquake.usgs.gov]. The black dashed lines show the locations of the 2.5-D gravity models.
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Figure 2. Simplified geologic map of the Central Andes showing major morphostructural units discussed in the text. Abbreviations: AT Altiplano, Pn Puna, FC Frontal Cordillera, CD Coastal Depression, CC Coastal Cordillera, FP Fore-arc Precordillera, WC Western Cordillera, EC Eastern Cordillera [after Charrier et al., 2013].
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Figure 3. Complete Bouguer anomaly map of the Central Andes from combined satellite and terrestrial gravity data. The gravity data are from the European Improved Gravity model of the Earth [EIGEN-6C2; Förste et al., 2012].
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Figure 4. Regional gravity anomaly map of the Central Andes.
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Figure 5. Residual gravity anomaly map of the Central Andes. The red triangles are for volcanoes.
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Figure 6. Residual topography of the Central Andes based on Airy isostasy model.
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(A)
(B)
Figure 7. 2.5-D gravity models of the Central Andes: (a) Profile at 16°S, (b) Profile at 19°S. 30
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Table 1. P-wave velocities and densities of major tectonic units in the Central Andes. The densities are derived from P-wave velocities using empirical relations [ Sobolev & Babeyko, 1994; Nafe and Drake, 1957].
Tectonic units
P- wave velocity (km s-1 )
Density (g cm-3 )
References
Continental domain: Upper & middle crust
5.0 - 7.0
2.54 - 2.97
Husen et al., 2000; Oncken et al., 2003; Krabbenhöft et al., 2004
Lower crust
6.7 -7.3
2.88 - 3.06
Husen et al., 2000; Oncken et al., 2003
Continental Mantle Lithosphere
7.6 - 8.1
3.2 - 3.33
Oncken et al., 2003
Oceanic domain: Sediment
2.0 – 4.5
1.91 – 2.5
Crust
6.5- 7.3
2.83 - 3.20
Husen et al., 200; Krabbenhöft et al., 2004 Krabbenhöft et al., 2004
Oceanic Mantle Lithosphere
7.8 - 8.15
3.22 – 3.35
Norabuena and Snoke, 1994; Krabbenhöft et al., 2004
Asthenosphere
8.15
3.35
Norabuena and Snoke, 1994
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Table 2. Density values of the final gravity model.
Tectonic units
Density (g cm-3 )
Variations (± g cm-3 )
Continental domain: Upper & middle crust Lower crust Continental Mantle Lithosphere
2.74 - 2.97 3.1 3.20 - 3.25
0.020 0.031 0.038
Oceanic domain: Sediment
2.55
0.106
Oceanic crust
2.90 - 3.20
0.025
Oceanic Mantle Lithosphere
3.20 - 3.36
0.034
Asthenosphere
3.35
0.023
33