Moho Depth Variations in the Taiwan Orogen from Joint Inversion of Seismic Arrival Time and Bouguer Gravity Data

Tectonophysics 632 (2014) 151–159

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Tectonophysics

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Moho depth variations in the Taiwan orogen from joint inversion of seismic arrival time and Bouguer gravity data

Zhiwei Li a,⁎, Steven Roecker b, Kwanghee Kim c,YaXud, Tianyao Hao d a State Key Laboratory of Geodesy and Earth's Dynamics, Institute of Geodesy and Geophysics, Chinese Academy of Sciences, Wuhan 430077, China b Department of Earth and Environmental Sciences, Rensselaer Polytechnic Institute, Troy, NY 12180, USA c Department of Geological Sciences, Pusan National University, Busan 609-735, Republic of Korea d Key Laboratory of Oil and Gas Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China article info abstract

Article history: The joint inversion of different geophysical datasets is an effective means to reduce the non-uniqueness and Received 4 March 2014 improve the reliability of geophysical inversion. In this study, seismic arrival time and Bouguer gravity datasets Received in revised form 3 June 2014 are jointly inverted to obtain an image of 3-D velocity structures in the Taiwan orogen. The model obtained Accepted 9 June 2014 from joint inversion fits the arrival time observations at least as well as when inverted individually, and the Available online 16 June 2014 gravity observations are much better fit when included in the inversion, implying a reduction in ambiguity by simultaneously modeling the disparate datasets. Moho depth variations estimated by the 3-D P wave velocity Keywords: Tomography model suggest a maximum Moho depth of 56 km located beneath the Backbone Central Range, and the trend Joint inversion of the Moho is largely consistent with the topography of the Central Range with eastward, asymmetric crustal Seismic arrival time thickening. The root beneath the Central Range appears to be smaller in lateral extent than previously imaged, Bouguer gravity anomaly and the velocity gradients into the uppermost mantle are significantly higher. The lack of evidence for a signifi- Subduction cant amount of Eurasian crust in the mantle supports geodynamical models of accreted, rather than consumed, Continental crust continental crust in a collisional environment. © 2014 Elsevier B.V. All rights reserved.

1. Introduction 1995; Roecker et al., 1987; Wang et al., 2006; Wu et al., 1997, 2007), Moho reflection and refraction phases (Hsu et al., 2011; Liang et al., The Taiwan orogen is the result of oblique collision between the 2007), receiver functions (Kim, et al., 2004; Wang et al., 2010a,b) and Eurasian plate (EAP) and the Philippine Sea plate (PSP) that active source seismic sounding (McIntosh et al., 2005, 2013; Wang commenced about 5 Ma ago (e.g., Wu et al., 1997)(Fig. 1). The fast et al., 2004) have been conducted to investigate the crustal structure be- convergence rate (~9 cm/yr) across the plate boundary between the neath the Taiwan orogeny. A key objective of these studies has been to EAP and PSP is responsible for the oceanic and continental plate determine the depth of the Moho, as variations in crustal thickness pro- collision and the significantly thickened crust of the Taiwan orogen vide important boundary conditions on mass balance at a convergent (e.g., Sella et al., 2002). The EAP subducts eastward under the PSP in boundary, including the ultimate fate (e.g., consumption vs. accretion) the south along the northern Luzon arc and the PSP subducts north- of Eurasian crust as it encounters the Luzon arc. To first order, the thick- westward under the EAP along the Ryukyu trench in the north of ening of the crust beneath the Taiwan orogeny, and the participation of Taiwan. This bidirectional subduction results in one of the most active the lower crust and upper mantle in the orogenic process, have been and complex orogens in the world. The high rate of seismicity and well established by both seismic arrival time tomography (Kim et al., dense seismological and geodetic networks make Taiwan an excellent 2005; Kuo-chen et al., 2012; Ma et al., 1996; Rau and Wu, 1995; Roecker place for investigating ongoing orogenic processes in an arc–continent et al., 1987; Wu et al., 1997, 2007) and 2-D and 3-D gravity modeling collision environment. Studies based on gravity anomalies (Hsieh (Hsieh et al., 2010; Mouthereau and Petit, 2003; Wang et al., 2004; et al., 2010; Yen et al., 1998), seismic arrival times (Kim et al., 2005; Yen et al., 1998; Zhang et al., 2004). However, the depth of the Moho be- Kuo-chen et al., 2012; Lin et al., 1998; Ma et al., 1996; Rau and Wu, neath Taiwan remains controversial and significant discrepancies exist from the analyses of different datasets using different techniques. While most studies agree that the crust is thickest beneath the Central Range, estimates of maximum Moho depth range from 33 km from ⁎ Corresponding author. Tel.: +86 27 86780717. gravity data inversion (Yen et al., 1998), 38 km from PmP traveltime in- E-mail address: [email protected] (Z. Li). version (Hsu et al., 2011), less than 54 km from receiver functions

Lishan fault

25˚

Taiwan Strait

nge

Western Foothills

24˚ Hsuehshan Range Hualien

rn Central Ra Backbone Range

Easte Ryukyu Trench

Coastal Range

Coastal Plain

23˚

it a ch tr ren Western Foothills n S T an yu

k Taitung rc u iw y Taiwa a R

c r

Longitudinal Valley ~9 on A on 0 mm/y

h uz L Philippine Sea

renc plate

la T

i Eurasian plate

Man

22˚ Hengchun A Luzon North Peninsula

120˚ 121˚ 122˚ 123˚

Fig. 1. Topographic map and the major tectonostratigraphic belts of Taiwan orogen. The Central Range is composed of the Backbone Range and the Eastern Central Range. Location of map is shown in the inset at the lower right.

(Wang et al., 2010a,b), 55 km from joint earthquake and active source Bouguer gravity data in the Taiwan orogen, and obtain a 3-D crustal tomography (Kuo-Chen et al., 2012), and 50–70 km from traveltime velocity model that can fit both sets of observations. tomography (Kim et al., 2005; Ma et al., 1996; Wu et al., 1997). Additionally, two Moho discontinuities for the EAP and PSP related 2. Method and data to the subduction polarity reversal were inferred beneath the Taiwan orogen from local earthquake tomography (Ustaszewski The travel times used in the seismic forward problem are calculated et al., 2012). using a finite difference eikonal equation solver in an earth-centered Because of differences in sampling and sensitivity, joint inversion spherical coordinate frame (e.g., Zhang et al., 2012), that has been of disparate geophysical datasets often can reduce model non- implemented previously in other tomographic studies (e.g., Kuo-chen uniqueness when compared to independent inversion (e.g., et al., 2012; Li et al., 2009, 2013; Roecker et al., 2004, 2006). Following Gallardo and Meju, 2011). Previous investigators have shown that Roecker et al. (2006), hypocenters are determined via a grid search ap- gravity data can provide additional constraints on the subsurface proach to ensure a global minimum in variance of arrival time residuals, structure when combined with seismic data. Examples include but are still included explicitly in the normal equations and determined joint inversion of gravity data with local earthquake arrival time simultaneously with velocity perturbations. The damped LSQR algo- data (Roecker et al., 2004; Roy et al., 2005; Wang et al., 2014), rithm (Paige and Saunders, 1982) is used for solving the system of linear teleseismic arrival time data (O'Donnell et al., 2011; Tiberi et al., equations. A coarse spherical mesh with approximate spacing of 6 km in 2003), active source seismic data (Tondi et al., 2000; Vermeesch horizontal and 4–20 km in vertical directions is used for specifying the et al., 2009), and surface wave dispersion curves (Maceira and model and accumulating partial derivatives, and an interpolated fine Ammon, 2009). Since Bouguer gravity data are sensitive mostly to mesh of ~2 km grid spacing is used for the finite difference traveltime density inhomogeneities of the crust and variations in the depth of calculation. Bouguer gravity observations are included in the inversion the Moho (e.g., Grad et al., 2009), a better subsurface model can be by presuming a functional relationship between P wave velocity and achieved by fitting both seismic arrival time and gravity observa- density. In this study we adopt the velocity–density relationships tions. In this study, we jointly invert seismic arrival time and determined by Gardner et al. (1974) and Christensen and Mooney Z. Li et al. / Tectonophysics 632 (2014) 151–159 153

(1995) for sedimentary, basement and mid-lower crustal rocks of con- (Fig. 2). To improve the reliability of the velocity model we use arrival tinental lithologies. Speciﬁcally, these are ρ = 9.89.3 + 289.1 ∗ VP for time data only from events that have been recorded at least by 6 stations 0.25 VP N 6.0 km/s, and ρ = 0.23 ∗ VP for VP b 5.5 km/s. A polynomial is and for which traveltime residuals are within ±3.0 s (approximately 3 used to obtain density in a smooth fashion between 5.5 and 6.0 km/s. times the standard deviation of the residuals). After selection, 568215 P The forward calculations of gravity and its partial derivatives are made and 443706 S arrivals from 45003 earthquakes are used in the tomo- directly in spherical coordinates with an adoptive numerical integration graphic inversion. Good ray path coverage is achieved in the crust and of the Gauss–Legendre quadrature (Asgharzadeh et al., 2007; Li et al., uppermost Moho beneath most of Taiwan. Bouguer gravity data used 2011). The strategy for simultaneous inversion of seismic arrival time in the joint inversion are taken from 603 readings with ~10 km spacing and gravity data follows that of Roecker et al. (2004). All observations or less (Yen et al., 1995), which cover most areas of the Taiwan orogen are included in each iteration, and a scalar factor is used to weight (Fig. 3). the relative sensitivities of arrival time and gravity data. An a posteriori moving window average smoothing of perturbations 3. Results with 3 × 3 × 3 grids is applied prior to model updates to mitigate possible artifacts appearing only at 1–2 grids. Finally, we assess the We conduct two independent inversions with the same parameters reliability and resolution of the results by conducting synthetic (i.e., model parameterization, starting model, damping, smoothing) in checkerboard tests. order to compare the images derived from arrival time only inversions

The seismic arrival time data used in this study were recorded by 98 with those from joint arrival time/gravity inversions. The starting VP permanent and temporary stations between 1991 and 2010. Most of the and VS models for both cases (Fig. 4)arebasedonone-dimensional arrival time picks are taken from International Seismological Center models used by Kim et al. (2005) and Wu et al. (2007),slightly (ISC), with some of them prior to 1998 provided by Prof. Futian Liu of adjusted to reduce arrival time residuals. In both inversions, no sig- the CAS Institute of Geology and Geophysics. As our focus is the crustal niﬁcant residual variance reduction was achieved after 7 iterations. structure beneath Taiwan, only earthquakes near land are selected For the arrival time only inversion, the root-mean-square (RMS) of

25˚

24˚

23˚

22˚

120˚ 121˚ 122˚ 123˚

Fig. 2. Distributions of seismic stations (red triangles) and earthquakes (closed circles). 154 Z. Li et al. / Tectonophysics 632 (2014) 151–159

Velocity (km/s) 123456789 0

25˚

20 Vs Vp

40 Depth (km) 24˚ 60

100 23˚

Fig. 4. The one-dimensional starting models for VP and VS used in both the joint arrival time/gravity inversion and the arrival time only inversion.

4. Discussion

Velocity models derived from arrival time/gravity data and from arrival time data only fit the arrival time observations equally well, but the forward gravity anomalies show significant discrepancies (Fig. 6). 22˚ Since the P wave velocity is related to the gravity observation through Bouguer Gravity the VP-density relations, differences in the P wave velocities are more pronounced than S wave velocities (Fig. 7). The main features of the P wave velocity models from joint inversion and arrival time only in- 120˚ 121˚ 122˚ version are virtually the same in the upper 40 km (Fig. 7), with differences increasing at depths greater than 50 km. While the patterns of velocity anomalies are similar below 50 km, the scale −90 -60 -30 0 30 60 90 and amplitude of the low P wave velocity anomalies beneath the Bouguer Gravity (mGal) Central Range are smaller and the gradient to the high P wave velocities east of the Taiwan orogen becomes larger in the arrival Fig. 3. Bouguer gravity anomaly data used in the joint inversion (Yen et al., 1995). time/gravity model. The primary reason for these differences is that these regions are relatively poorly sampled by the arrival time data. Hence these velocities can change without much effect on the arrival arrival time residuals decreased from 0.73 s to 0.43 s for P arrivals time misfit. and from 0.89 s to 0.50 s for S arrivals. For the joint arrival time/ Comparing the arrival time only and joint arrival time/gravity im- gravity inversion, the RMS of arrival time residuals decreased from ages (Fig. 7), it is clear that the main advantage of including Bouguer 0.73 s to 0.42 s for P arrivals and from 0.89 s to 0.50 s for S arrivals. gravity data in a joint inversion is that it significantly improves the def- The similarity in misfit shows that both models explain the seismic inition of the Moho, particularly at depths greater than about 40 km. arrival time data equally well. However, while the RMS of the gravity Choosing a contouring surface of VP = 7.5 km/s as a means to estimate residuals increased from 31.6 mGal to 34.8 mGal for the arrival time the Moho depth beneath the Taiwan orogen (Kuo-Chen et al., 2012), the only model, it decreased from 31.6 mGal to 13.0 mGal for the arrival main difference is that the root under the Central Range is narrower, time/gravity model. and the gradients in the upper mantle sharper, than the relatively We evaluate the resolution of the images using checkerboard tests deeper and broader root in the arrival time model. Note that the arrival with grid spacing of 6 km and approximate anomaly sizes of 12 km, time only model suggests the presence of low wavespeed material 18 km, 24 km and 30 km (Fig. 5). The patterns and amplitudes of check- extending both to the base of the model at 60 km depth and to the erboard models are well recovered under most of Taiwan at depths less east at depths between 40 and 60 km depth. One might infer from than about 50 km depth, with the best recovery in the middle to north- this image the subduction of the Eurasian continental crust as a means ern Central Range. At depths greater than about 60 km, the anomalies to obtain these low wavespeed anomalies. The joint inversion model, near the Central Range are well recovered, largely because of the ray however, shows that this extended low wavespeed zone is largely an ar- paths from the many deep earthquakes in this area. tifact that results in the misfit in the gravity observations along the east Z. Li et al. / Tectonophysics 632 (2014) 151–159 155

Depth:20 km Depth:50 km 25˚00' Vp 25˚00' Vp

24˚30' 24˚30'

24˚00' 24˚00'

23˚30' 23˚30'

23˚00' 23˚00'

22˚30' 22˚30'

22˚00' ~24x24km 22˚00' ~30x30km

120˚00' 120˚30' 121˚00' 121˚30' 122˚00' 120˚00' 120˚30' 121˚00' 121˚30' 122˚00'

6.2 6.4 6.6 7.6 7.8 8.0 8.2 Vp (km/s) Vp (km/s)

Fig. 5. Representative results from the checkerboard tests of resolution at 20 km and 50 km depths. A larger scale anomaly is taken at 50 km because of the reduced level of resolution at this depth.

coast of Taiwan (Fig. 6). This joint inversion model therefore argues polarity (Ustaszewski et al., 2012). Such complications could make the against the subduction of significant amounts of Eurasian crust into identification of interfaces producing mode conversions problematic. the mantle, as has been suggested for continental collisions in central Finally, the Moho map determined by joint inversion is similar to that Asia (e.g., Roecker, 1982; Schneider et al., 2013; Sippl et al., 2013), but obtained by Zhang et al. (2004), who carried out a crustal structure in- supports geodynamical models for crustal accretion. version using gravity data with a 3-D seismic velocity model as an initial The main features of velocity structure that we find in the upper model, but it is not clear how compatible their model is with the seismic 40 km are similar to those reported in previous local earthquake tomog- observations. raphy studies (e.g., Kim et al., 2005; Wu et al., 2007). These include the low-VP,low-VS and high-VP/VS anomalies found in the Coastal Plain at 0 5. Conclusions and 4 km depth, which extend eastward to the Western Foothills at 8 and 12 km depth, and the low-VP,low-VS and normal-to-low VP/VS be- We have carried out a joint inversion of seismic arrival time and neath the Central Range that become more predominant at 16 to gravity data in the Taiwan orogen. The 3-D velocity model obtained 40 km depth. from the joint inversion fits both seismic arrival time and gravity Previous studies also have suggested that the crust has thickened to observations at approximately the noise level of the observations. more than 50 km beneath the Central Range (e.g., Kim et al., 2005; Rau Comparisons between the models from joint inversion and arrival and Wu, 1995; Wu et al., 2007). The asymmetrically thickened crust time only inversion show that the well resolved velocity structures shown clearly in the near E–Wprofiles (Fig. 7a–d) is similar to that remain almost the same for regions with good ray path coverage. obtained from joint local and teleseismic tomographic results (e.g., Gravity data provide additional constraints on the velocity structure Kuo-Chen et al., 2012). The Moho depth from the arrival time/gravity near the Moho and in the uppermost mantle where seismic ray paths model (Fig. 8) is close to the estimated Moho depth obtained by Kuo- are relatively sparse. Hence, the reliability of the Moho depth esti- Chen et al. (2012) using the denser seismic stations and active sources mated from the velocity model is improved by the joint inversion of TAIGER project, although we note that, as with our arrival time only of seismic arrival time and Bouguer gravity data. The main result model, their model does not fit the Bouguer gravity data particularly from the image generated by this study is that the root under the well (their model misfits the east–west trend in gravity by a total of Central Range is narrower, and the gradients in the upper mantle about 100 mGal). Comparison with the crustal thickness results from sharper than previously determined. This model therefore favors H–κ stacks of receiver functions (Wang et al., 2010a,b)givesMoho the accretion of Eurasian crust, as opposed to subduction of signifi- depths within a few kilometers for most stations in the central part of cant amounts of lower crust into the mantle. the Taiwan orogeny (Fig. 8). Note that the large differences are located mostly in the eastern and northern margin of Taiwan, where the crust Acknowledgments and upper mantle appear to be complicated by the history of subduction in this area — for example, two Moho discontinuities of the EAP and PSP We thank the editor, Prof. Laurent Jolivet, and two anonymous re- were proposed to exist here as a result of the change in subduction viewers for their valuable comments. We thank Prof. Sidao Ni for 156 Z. Li et al. / Tectonophysics 632 (2014) 151–159

(a) (b) 25˚ 25˚ Calculated Residual

24˚ 24˚ A

23˚ 23˚ A’

Joint Joint 22˚ Inversion 22˚ Inversion

120˚ 121˚ 122˚ 120˚ 121˚ 122˚

24˚ 24˚

23˚ 23˚

Arrival Arrival 22˚ time only 22˚ time only

120˚ 121˚ 122˚ 120˚ 121˚ 122˚

−90 −60 −30 0 30 60 90 Bouguer Gravity (mGal)

Fig. 6. (a) Calculated gravity anomaly data from the joint arrival time/gravity model and (b) associated residuals from the Bouguer gravity observations; (c) calculated gravity anomaly data from arrival time only model and (d) associated residuals from the Bouguer gravity observations.

helpful discussions. This work was supported by NSFC41210005, References NSFC41304045, XDB06030203 and SKLGED2013-9-3-Z. Participa- tion of S. Roecker in this study was supported by NSF Grant EAR- Asgharzadeh, M.F., von Frese, R.R.B., Kim, H.R., Leftwich, T.E., Kim, J.W., 2007. Spherical prism gravity effects by Gauss–Legendre quadrature. Geophys. J. Int. 169, 1–11. 1010580. H. Kim in this study was funded by the KMA, Research De- Christensen, N.I., Mooney, W.D., 1995. Seismic velocity structure and composition of the velopment Program under Grant CATER 2014-5030. continental crust: a global view. J. Geophys. Res. 100, 9761–9788. Z. Li et al. / Tectonophysics 632 (2014) 151–159 157

AA’AA’Distance (km) Distance (km) 0 50 100 150 200 0 50 100 150 200 3 3 0 0 −3 −3 0 0 Elev (km) Elev (km) 10 (a) 10 (b) 20 20 6.5 6.5 7 30 7.5 30 7 7.5 40 40 50 50 Depth (km) 60 Depth (km) 60 70 Joint Inversion Vp 70 Arrival time only Vp 80 80 120.0 120.5 121.0 121.5 122.0 120.0 120.5 121.0 121.5 122.0 Longitude (deg) Longitude (deg)

5.0 5.5 6.0 6.5 7.0 7.5 8.0 5.0 5.5 6.0 6.5 7.0 7.5 8.0 Vp (km/s) Vp (km/s) Distance (km) Distance (km) 0 50 100 150 200 0 50 100 150 200 3 3 0 0 −3 −3 0 0 Elev (km) Elev (km) 10 (c) 10 (d) 20 3.75 20 3.75 4 4 30 4.25 30 4.25 40 40 50 50 Depth (km) 60 Depth (km) 60 70 Joint Inversion Vs 70 Arrival time only Vs 80 80 120.0 120.5 121.0 121.5 122.0 120.0 120.5 121.0 121.5 122.0 Longitude (deg) Longitude (deg)

2.8 3.0 3.2 3.4 3.6 3.8 4.0 4.2 4.4 4.6 2.8 3.0 3.2 3.4 3.6 3.8 4.0 4.2 4.4 4.6 Vs (km/s) Vs (km/s) Distance (km) Distance (km) 0 50 100 150 200 0 50 100 150 200 3 3 0 0 −3 −3 0 0 Elev (km) Elev (km) 10 10 (e) 1.8 (f) 20 20 30 30 1.7 40 1.7 40 50 50 Depth (km) 60 Depth (km) 60 70 Joint Inversion Vp/Vs 70 Arrival time only Vp/Vs 80 80 120.0 120.5 121.0 121.5 122.0 120.0 120.5 121.0 121.5 122.0 Longitude (deg) Longitude (deg)

1.60 1.65 1.70 1.75 1.80 1.85 1.60 1.65 1.70 1.75 1.80 1.85 Vp/Vs Vp/Vs

Fig. 7. (a) and (b) show the P wave velocity variations along AA’ proﬁle (shown in Fig. 6b) for the joint arrival time/gravity model and the arrival time only model, respectively; (c) and (d) are the same as (a) and (b) for S wave velocity; (e) and (f) are the same as (a) and (b) but for VP/VS. Note that structures off of the eastern coast of Taiwan (i.e., distances greater than about 150 km) generally are not well constrained by this inversion.

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(a) (b) −14.9±1.6 −15.4±1.6 −10.1±2.5 25˚ Vp=7.50 km/s Vp=7.50 km/s −9.1±2.5 −6.5 25˚ −5.5 −17.5 −17.4 −0.2 −11.6 −0.3 −11.2 3.6±1.2 402.1±1.2 −4.2 −4.3

−14.1±3.1 −18±3.1 −1.6±1.8 −2.2±1.8

24˚ 40 −14.7 24˚ 40 −15.7 −5.8±6.0 −5.5±6.0 40 40 1.2 1.6 50 50 −4.3±3.4 −10±3.4 −3.6±2.0 −5.2±2.0

4.7 3.4 23˚ 23˚ −2±1.7 −0.9±1.7 30

3.6 −7.1 4 −4.6

22˚ Joint inversion 22˚ Arrival time only

120˚ 121˚ 122˚ 120˚ 121˚ 122˚

25 30 35 40 45 50 55 60 Depth (km)

Fig. 8. Contour surface of VP = 7.5 km/s from the (a) joint arrival time/gravity model and (b) arrival time only model, taken as a proxy for the depth to Moho. Colored circles summarizethe Moho depth derived from H–κ staking of receiver functions (Wang et al., 2010a,b). White numbers indicate the difference in km between the receiver function and joint inversion estimates, along with the associated uncertainties (where available) in the receiver function estimates. Note that the area near N23.75°, E120.5° is masked by gray in the joint inversion result because the difference in models results in this area being sampled by less than 6 rays.

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