Crustal structure across the NE Tibetan Plateau and Ordos Block from

Tectonophysics 705 (2017) 33–41

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Crustal structure across the NE Tibetan Plateau and Ordos Block from the joint inversion of receiver functions and Rayleigh-wave dispersions Yonghua Li a,b,⁎, Xingchen Wang a,b, Ruiqing Zhang b, Qingju Wu a,b, Zhifeng Ding a,b a b

Key Laboratory of Seismic Observation and Geophysical Imaging, China Earthquake Administration, Beijing 100081, China Institute of Geophysics, China Earthquake Administration, Beijing 100081, China

a r t i c l e

i n f o

Article history: Received 20 October 2016 Received in revised form 4 February 2017 Accepted 24 March 2017 Available online 27 March 2017 Keywords: Crustal structure Northeastern Tibetan Plateau Ordos Block Rayleigh-wave dispersion Receiver function Joint inversion

a b s t r a c t We investigated the crustal structure at 34 stations using the H-κ stacking method and jointly inverting receiver functions with Rayleigh-wave phase and group velocities. These seismic stations are distributed along a profile extending across the Songpan-Ganzi Terrane, Qinling-Qilian terranes and southwestern Ordos Basin. Our results reveal the variation in crustal thickness across this profile. We found thick crust beneath the Songpan-Ganzi Terrane (47–59 km) that decreases to ~45–47 km in the west Qinling and Qilian terranes, and reaches its local minimum beneath the southwestern Ordos Block (43–51 km) at an average crustal thickness of 46.7 ± 2.5 km. A low-velocity zone in the upper crust was found beneath most of the stations in NE Tibet, which may be indicative of partial melt or a weak detachment layer. Our observations of low to moderate Vp/Vs (1.67–1.79) represent a felsic to intermediate crustal composition. The shear velocity models estimated from joint inversions also reveal substantial lateral variations in velocity beneath the profile, which is mainly reflected in the lower crustal velocities. For the Ordos Block, the average shear wave velocities below 20 km are ~3.8 km/s, indicating an intermediate-to-felsic lower crust. The thick NE Tibet crust is characterized by slow shear wave velocities (3.3– 3.6 km/s) below 20 km and lacks high-velocity material (Vs ≥ 4.0 km/s) in the lower crust, which may be attributed to mafic lower crustal delamination or/and the thickening of the upper and middle crust. © 2017 Elsevier B.V. All rights reserved.

1. Introduction The NE margin of the Tibetan Plateau is recognized as one of ideal sites to study continental dynamics and plateau building. This region is tectonically complex, with a variety of crust and fault types and strong seismicity (Fig. 1). The present tectonics of NE Tibet resulted from SW– NE compression induced by collision between India and Eurasia since ~50 Ma (Taylor and Yinn, 2009; Yin and Harrison, 2000). Recent Global Positioning System (GPS) measurements (Gan et al., 2007) show that about 10 mm/yr of NE–SW or NNE–SSW horizontal shortening is being accommodated in NE Tibet. Geological surveys (Meyer et al., 1998) showed that active crustal shortening occurs prominently along and north of the Kunlung Range. Numerous passive and active seismic studies (e.g., Li et al., 2006, 2014b; Liu et al., 2006, 2014; Pan and Niu, 2011; Teng et al., 2013; Tian and Zhang, 2013; Vergne et al., 2002; Wang et al., 2016; Xu et al., 2014; Zhang et al., 2011; Zheng et al., 2016) showed that the crust thickens from ~42 km beneath the Ordos Basin to ~60 km beneath the Qinling-Qilian terranes and ~ 70 km beneath the Songpan-Ganzi Terrane. However, mechanisms that try to explain crustal shorting and ⁎ Corresponding author at: Key Laboratory of Seismic Observation and Geophysical Imaging, China Earthquake Administration, Beijing 100081, China. E-mail addresses: [email protected], [email protected] (Y. Li).

http://dx.doi.org/10.1016/j.tecto.2017.03.020 0040-1951/© 2017 Elsevier B.V. All rights reserved.

thickening under NE Tibet are controversial. For example, a widespread midcrustal low-velocity zone(LVZ) beneath the Songpan-Ganzi Terrane and Kunlun Mountains, revealed by recent surface wave studies based on ambient noise and earthquake data (Bao et al., 2013; Jiang et al., 2014; Li et al., 2014a; Yang et al., 2012), supports the crustal flow model of Clark and Royden (2000). This model predicted that the middle-lower crust would flow northward from central Tibet and crustal thickening in the NE Tibetan Plateau would occur mainly in the lower crust. However, the low crustal Vp/Vs ratios across the NE Tibetan Plateau, demonstrated by previous receiver function studies (Li et al., 2006; Pan and Niu, 2011; Tian and Zhang, 2013; Vergne et al., 2002; Wang et al., 2016), have been used to suggest that thickening of the NE Tibetan crust is predominantly caused by upper crustal shortening (Li et al., 2006; Tian and Zhang, 2013; Vergne et al., 2002) or vertically coherent shortening (Pan and Niu, 2011; Wang et al., 2016). The Ordos Plateau, which belongs to the western part of the SinoKorea Craton, is an important boundary of NE Tibet and has a Paleoproterozoic basement (Wan et al., 2013). Investigations of passive and active source seismic detection have been conducted across the Ordos Plateau, but there is still debate regarding the crustal thickness and composition of the lower crust across this block. For example, receiver function and active source seismic refraction/reflection studies suggest that crustal thickness ranges from 40 to 50 km beneath the Ordos Basin (Jia et al., 2014; Liu et al., 2006; Wang et al., 2014, 2016).

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Y. Li et al. / Tectonophysics 705 (2017) 33–41

Fig. 1. Tectonic map of the study area showing topography and tectonic features (Taylor and Yinn, 2009). The inset map shows the location of the study area. Suture zones are marked by white dashed lines. AMS: Anyimaqen-Kunlun-Muztagh suture zone. Red solid lines denote faults. HYF: Haiyuan fault; KF: Kunlun fault; WQLF: Western Qinling fault; LMT: Longmenshan fault; LPSF: Liupanshan fault. The seismic stations (solid triangles) used in the study are distributed along a seismic profile line (A–B). The red triangles denote the locations referred to in Figs. 3 and 4.

Based on the P-wave velocity and Passion ratios, Liu et al. (2006) proposed that the lower crust under the Ordos Block has an intermediateto-felsic lithology. Virtual deep seismic sounding and receiver function analyses by Yu et al. (2012), however, concluded that the interface at 40 km is the Conrad discontinuity, and suggested that the lower crust may contain a thick (~20 km) layer of mafic rock. Determining the lateral variation in crustal structure across NE Tibet and the Ordos Block is important for improving our understanding of crustal growth and tectonic deformation mechanisms in this region. In this study, we computed receiver functions using data recorded by a linear array of 34 broadband seismic stations from the ChinArray (Phase II) project (Fig. 1). We report new estimates of crustal structure (thickness, shear wave velocity and Poisson's ratio) along a profile extending across NE Tibet and the Ordos Basin by using the H-κ stacking method and a joint inversion of Rayleigh-wave dispersions and P-wave receiver functions. The resulting crustal structures provide valuable insight into tectonic evolution processes in this region. 2. Data and methods

distance of 35 km. Each station consisted of a Guralp CMG-3ESPC or CMG-3ESP seismometer and a Reftek 130 data acquisition system sampling at 100 samples per second (sps). The 34 broadband seismic stations are deployed along a profile extending from the Songpan-Ganzi Terrane, through the Qinling-Qilian terranes, to the southwestern Ordos Basin (Fig. 1). Nearly two years of data from these stations were used for this study. Rayleigh-wave phase velocities at periods ranging from 10 s to 80 s and group velocities for periods from 10 s to 60 s from a surface tomographic study (Li et al., under review) were used for the joint inversion. Rayleigh-wave phase and group velocity dispersions were based on an analysis of teleseismic waveform data recorded by 650 seismic stations from ChinArray (Phase II). More than 18,000 inter-station phase and group velocity dispersion curves were estimated using the two-station method. The obtained dispersion curves were then tomographically inverted using the 2-D tomography method of Barmin et al. (2001) to construct 2-D phase and group velocity maps on a 0.5° × 0.5° grid covering the study region. The tests presented by Li et al. (under review) showed that the resolution length of the tomographic maps is mainly b100 km in the study area.

2.1. Data Three-component seismograms, recorded at 34 stations from ChinArray (Phase II), were used in this study to compute receiver functions. The ChinArray (Phase II) began operation since September 2013 and consists of 650 portable seismic stations spread across the NE margin of the Tibetan Plateau and its adjacent areas, with an average station

2.2. Calculation of P-wave receiver functions P-wave receiver functions were computed using seismograms from 169 teleseismic events with epicenter distances from 30° to 90° and magnitudes N5.5 (Fig. 2).


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crustal thickness as well as Vp/Vs values are given in Table 1. Receiver functions with higher frequencies (α = 2.0) were used for all stations except 51503, 62346, 62373 and 62400, where lower frequency (α = 1.0) receiver functions were used. For these stations, the coherent Moho Ps phases and multiples were more clearly identified at lower frequencies than at higher frequencies. Results from two stations (62363 and 51551) were not included in this study because of a lack of sufficient receiver function waveforms. The uncertainties in H and κ were assessed using bootstrapping resampling with 100 samples from the original receiver function datasets. The uncertainties for each station, using an average crustal Vp of 6.3 km/s for the H-κ stacking, are shown in Table 1. To evaluate the effect of varying mean crustal Vp on the estimation of H and κ, the H-κ stacking was recomputed using P-wave velocities of 6.15 and 6.45 km/s (Supplementary Fig. A.1). The combination of the two methods shows that the overall uncertainties are ±2–3 km for crustal thickness and ±0.05 for Vp/Vs at each station. 2.4. Joint inversion of receiver functions and Rayleigh-wave dispersion Fig. 2. Location of teleseismic events (small blue circles) used in receiver function analysis. The red triangle is the center of the profile. The numbers on the larger circles show the epicenter distance in degrees.

The selected seismograms were decimated to 10 sps, windowed between 10 s before and 100 s after the theoretical P-wave onset and band pass filtered (0.05–4 Hz) to remove low-frequency noise. We rotated two horizontal component seismograms to obtain radial and transverse components. Receiver function was obtained by deconvolving the vertical component with the radial and tangential components. In this study, two overlapping frequency bands using Gaussian filter parameters of α = 1.0 (f ≤ 0.5 Hz) and α = 2.0 (f ≤ 1.0 Hz) were used to compute the receiver functions using the time domain iterative deconvolution method of Ligorria and Ammon (1999). Coherent receiver functions were automatically selected by using cross-correlation coefficients of radial receiver functions. The remaining coherent receiver function waveforms were further visually checked and a total of 2300 high-quality receiver functions for each frequency were retained at 34 stations. Fig. 3 shows the receiver function with Gaussian width factor 1.0 for stations 61063 and 62335 (Fig. 3c, d).

2.3. H-κ stacking analyses The H-κ stacking technique of Zhu and Kanamori (2000) was applied to the receiver functions to estimate the crustal thickness (H) and mean Vp/Vs (κ) beneath each station. The stacking procedure compares the observed amplitudes and arrival times of Moho Ps converted phases and their multiples (PpPs, PpSs + PsPs) with predictions for ranges of H and Vp/Vs. The best estimations of crustal thickness and Vp/Vs are obtained when the three phases are stacked coherently. We applied stacking technique by searching over broad ranges of crustal thickness (35–70 km) and Vp/Vs (1.50–2.00). In applying the stacking procedure, it is necessary to choose a mean crustal Vp and weights for the converted phases. An average Vp was set to 6.3 km/s, according to previous wide angle reflection/refraction studies (Liu et al., 2006; Wang et al., 2003, 2007; Zhang et al., 2008). For all stations, the weights 0.6, 0.3, and 0.1 were assigned to Ps, PpPs and PpSs + PsPs, respectively. Two examples for stations 61063 and 62335 are shown in Fig. 3 with predicted arrival times of Moho Ps and the two multiples for the optimal H and k values (Fig. 3c, d). In this study, the H-κ stacking technique was performed on receiver functions with two overlapping frequency bands; similar estimates were obtained for each. Results from the H-κ stacking method for

The 1-D shear wave velocity models for each site were obtained by jointly inverting the receiver functions with fundamental mode Rayleigh-wave phase and group velocities. The advantage of simultaneously using receiver functions and surface wave dispersion velocities is that tighter constraints on the Vs structure can be obtained compared with modeling each data set independently (Julià et al., 2000, 2005). The joint inversion method used in this study is an iterative damped least squares technique (Julià et al., 2000; Herrmann, 2013), which has been used in numerous investigations (e.g., Julià et al., 2000, 2005; Zheng et al., 2016). In the joint inversion, the receiver functions for each station are divided into four groups with different ray parameter bins (0.04–0.049, 0.05–0.059, 0.060–0.069, and 0.070–0.079). For each group of receiver functions, the average receiver functions with high- and low-frequency content were computed. These averaged receiver functions were obtained by stacking all events for a broader range of back azimuths. The number of receiver functions used to create these stacks depends on the station, but, in general, is N 10 for each group. Previous studies showed that there are variations in receiver functions for different azimuths at some stations within northeastern Tibet (Wang et al., 2016). These small variations in the receiver functions may be related to the presence of anisotropy and/or dipping layers. However, we found that these differences are quite small and the Moho Ps phase is generally identical for different azimuths. Thus, the resulting structure from inverted average receiver functions reflects a firstorder tectonic characteristic beneath a given station. For each station, Rayleigh-wave dispersion curves used for our joint inversion were extracted from the corresponding tomographic grids containing that station. The initial velocity model used in the joint inversion has a layer thickness of 1 km at the top of the model, 2-km thickness between 3and 100-km depth, 5 km between 100- and 150-km depth, and 10 km below a depth of 150 km. The model was parametrized with a constant shear velocity of 4.48 km/s down to 100 km, overlying a flattened Ak135 model (Kennett et al., 1995). The starting model with a layer velocity of 4.48 km/s, which is the upper mantle velocity in the Ak135 model (Kennett et al., 1995), ensure that no a priori assumptions about the location of the Moho are made. The crustal Vp/Vs values in each layer is given according to the receiver function H-κ result for each station. Crustal densities were computed from the P-wave velocities through the empirical relationship of Berteussen (1977). A differential smoothing constraint was applied to find the best model that minimizes a weighted combination of least squares norms for each data set. To stabilize the inversion, a slightly higher damping value of 10 was adopted in the initial two iterations and then a damping value of 0.1 was chosen for the following iterations. Another important

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Fig. 3. H-κ stacking analysis of receiver functions at station 61063 (a, c) and 62335 (b, d). The location of the two stations is shown in Fig. 1. (a, b) H-κ stacking results. Preferred H-κ values (white circles) and the corresponding error (ellipse) are marked. (c, d) Radial receiver functions (Gaussian filter of 1.0) plotted according to the ray parameter. The theoretical arrivals (black dashed lines) of the Ps, PpPs and PpSs + PsPs phases are calculated using the optimum values for H and k.

parameter is the influence factor, p (0 ≤ p ≤ 1), which is used to adjust the relative weight of receiver functions and dispersion data in the inversion. Setting p equal to 0 or 1 result in a solution only based on receiver function or surface wave dispersion, respectively. We tested the influence of different p values (p = 0.3, 0.5 and 0.7) and selected p = 0.3 to give more weight to the receiver function and define a shear wave velocity model. Two examples illustrating the determination of the shear wave velocity model are provided in Fig. 4. One station located on the NE Tibetan Plateau with low-velocity lower crust and another on the Ordos Block with a relatively higher-velocity lower crust. The fit to the receiver function stacks and the surface wave dispersions is excellent, although some high-frequency signals in the receiver functions are not fully modeled (Fig. 4). Uncertainties of the resulting shear wave velocity models were estimated following the approach of Julià et al. (2005) by repeatedly performing the inversions using a range of influence parameters (0.1, 0.3, and 0.5) and selectively discarding receiver function and surface wave observations. Using this method, we repeatedly performed the inversions for two selected stations. The resulting 100 inversions for each station show that the error of the shear wave velocity in each layer is about 0.1–0.2 km/s and the uncertainties of crustal thickness is about ±2 km (Supplementary Fig. A.2). Previous studies showed that shear wave velocity in the lower crust are b4.3 km/s and shear wave velocities N4.3 km/s are typical for mantle lithologies (Christensen, 1996; Christensen and Mooney, 1995). Therefore, estimates of crustal thicknesses in the joint inversions are based on where the shear velocity jump is larger and where velocity is close to or exceeds 4.3 km/s. The inferred crustal thicknesses are shown in Table 1. Following the approach used in previous studies (e.g., Julià et al., 2005), we also estimated the thickness of the mafic

lower crust (shear wave velocities ≥ 4.0 km/s), which is indicative of the sharpness of Moho. 3. Results The shear wave velocity model for the 34 stations was determined by joint inversion. A 2-D shear wave velocity model along a seismic profile (AB in Fig. 1) was constructed by combining individual 1-D velocitydepth profiles obtained for each station. The 2-D velocity profile across NE Tibet and the Ordos Basin is shown in Fig. 5 that displays the velocity distributions with depths up to 70 km. Our results for crustal thickness, Poisson's ratio, lower crustal shear wave velocity, mafic lower crust thickness, and average crustal Vs are summarized in Table 1. The overall uncertainty in Moho depth for each station is ~ 2–3 km. The crustal thicknesses determined from the joint inversions are in good agreement with those obtained from the H-κ stacking method for all stations (Fig. 5 and Table 1); thus, we use the crustal thickness obtained from the joint inversion for our interpretation. Our investigation reveals significant crustal thickness variations along the profile (Fig. 5 and Table 1). The crustal thickness increases gradually from 43 km beneath the Ordos Basin in the northeast to 59 km beneath the Songpan-Ganzi Terrane in the southwest. The crustal thickness in the Songpan-Ganzi Terrane is quite variable, ranging from 59 to 47 km. The average crustal thickness beneath this area is 52.8 ± 2.9 km, which is the largest value of the entire profile. The average crustal thickness is 45.6 ± 1.0 km beneath the West Qinling Orogen, 47.0 ± 1.6 km beneath the Qilian Orogen, and 46.7 ± 2.6 km beneath the Ordos Basin. Our results are consistent with most of the previous estimations of crustal thickness using receiver function stacking (Li et al., 2014b;


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Table 1 Crustal structure results at the 34 seismic stations used in this study. Tetonic region

Station name

Lat (°)

Lon (°)

Elevation (m)

Na

Dep1b (km)

Vp/Vs

Dep2c (km)

Average c rustal Vs (km/s)

Average crustal Vs below 20 km (km/s)

Mafic lower crust (km)

Songpan-Ganzi Terrane

51563 51553 51552 51555 51541 51538 51551 51537 51536 51501 51503 62312 Mean 62323 62319 62328 62324 62335 62342 Mean 62336 62346 62363 62354 Mean 62358 62373 62380 62389 62385 62400 62403 61062 61063 61067 61064 61070 Mean

31.85 32.17 32.21 32.49 32.25 32.63 33.01 32.92 32.99 33.56 33.51 33.65

101.70 102.16 102.49 102.44 102.88 103.09 102.98 103.39 103.69 103.67 104.01 104.43

2914 2749 3586 3590 2974 3226 3519 3462 3412 2503 1918 1903

101 74 57 97 62 91 8 100 57 57 89 33

34.05 33.92 34.18 34.12 34.45 34.71

104.39 104.88 104.91 105.15 104.96 105.25

1767 2125 1784 1658 2236 1587

61 44 66 69 82 103

34.47 34.84 35.27 35.14

105.55 105.80 105.96 106.21

1733 1480 1828 1888

15 54 52 76

35.18 35.60 35.78 36.05 36.01 36.36 36.56 36.64 36.68 36.96 36.85 37.17

106.57 106.64 107.02 107.04 107.49 107.51 107.89 108.10 108.47 108.38 108.78 108.77

1612 1667 1272 1268 1399 1449 1445 1416 1212 1438 1436 1602

50 46 33 38 40 83 64 98 98 95 124 92

59.0 ± 1.5 54.9 ± 0.4 53.0 ± 1.1 52.9 ± 1.1 52.7 ± 1.1 52.0 ± 0.6 – 51.9 ± 0.7 54.0 ± 1.8 49.1 ± 1.2 48.4 ± 1.2 45.9 ± 1.2 52.1 44.6 ± 0.9 42.0 ± 0.7 42.9 ± 0.8 43.0 ± 1.6 44.6 ± 0.9 44.1 ± 1.1 43.5 – 44.8 ± 1.5 44.9 ± 1.2 45.5 ± 1.1 45.0 48.5 ± 0.7 51.9 ± 1.2 49.6 ± 1.7 49.5 ± 1.2 49.3 ± 0.9 46.8 ± 1.5 46.5 ± 0.8 47.5 ± 0.6 46.4 ± 0.5 43.4 ± 0.9 44.0 ± 1.1 42.5 ± 2.1 47.2

1.71 ± 0.03 1.70 ± 0.02 1.79 ± 0.03 1.75 ± 0.03 1.75 ± 0.03 1.76 ± 0.02 – 1.78 ± 0.02 1.77 ± 0.08 1.79 ± 0.04 1.71 ± 0.04 1.72 ± 0.04 1.75 1.75 ± 0.04 1.76 ± 0.02 1.75 ± 0.04 1.76 ± 0.04 1.71 ± 0.03 1.71 ± 0.03 1.74 – 1.69 ± 0.04 1.79 ± 0.03 1.78 ± 0.03 1.75 1.78 ± 0.05 1.67 ± 0.03 1.75 ± 0.04 1.74 ± 0.04 1.71 ± 0.03 1.78 ± 0.04 1.76 ± 0.03 1.72 ± 0.02 1.71 ± 0.01 1.73 ± 0.04 1.76 ± 0.05 1.77 ± 0.06 1.74

59 55 55 53 53 53 53 53 51 51 51 47 53 45 45 47 45 47 45 46 45 47 49 47 47 49 47 51 49 49 47 47 47 45 43 43 43 47

3.5 3.5 3.5 3.4 3.5 3.5 3.5 3.5 3.5 3.5 3.6 3.6 3.5 3.5 3.6 3.6 3.6 3.6 3.6 3.6 3.6 3.6 3.6 3.6 3.6 3.7 3.7 3.8 3.7 3.7 3.8 3.7 3.7 3.7 3.7 3.6 3.7 3.7

3.6 3.6 3.6 3.5 3.5 3.5 3.5 3.6 3.6 3.5 3.6 3.7 3.6 3.6 3.7 3.7 3.7 3.7 3.7 3.7 3.7 3.8 3.8 3.7 3.7 3.8 3.8 3.9 3.8 3.9 3.8 3.8 3.8 3.8 3.7 3.7 3.7 3.8

4 4 4 4 4 2 4 2 2 2 4 2 3 2 4 4 4 4 4 4 2 4 2 2 3 4 4 4 4 6 2 6 4 2 2 2 2 4

West Qinling Terrane

Qilian Terrane

Ordos Block

a b c

N, number of receiver function waveforms. Dep1, crustal thickness from H-k stacking technology. Dep2, crustal thickness from joint inversion.

Pan and Niu, 2011; Wang et al., 2016; Zheng et al., 2016) and with results from seismic refraction/reflection data (Liu et al., 2006; Wang et al., 2003; Zhang et al., 2008). Wang et al. (2014) used the receiver function stacking technique and proposed that the average crustal thickness beneath the southern part of the Ordos Basin is 39.5 km, which is thinner than our results. Despite this discrepancy, their observations in other terranes are in general agreement with our results. The resulting Vp/Vs values of all stations range from 1.67 to 1.79 with an average of 1.74 ± 0.03, which is in agreement with most of the previous estimations of crustal Vp/Vs ratios (Wang et al., 2014; Zheng et al., 2016). In addition, we found little differences in the average Vp/Vs between different terranes along this profile (Table 1). Using data from the CEA regional network, Wang et al. (2016) suggested that the average Vp/Vs values beneath the Ordos Basin (1.794 ± 0.024) are considerably larger than the remaining terranes (b1.72). Those observations are slightly different from our results, but are still comparable within the reported uncertainties. Our results show that crustal velocity structure varies considerably along this profile. Almost all of the stations located within the Songpan-Ganzi Terrane, West Qinling Orogen and southern part of the Qilian Orogen display a strong increase of S-wave velocities from 3.0 to 3.6 km/s down to a depth of about 10 km. In the 12–20-km-depth interval, lower velocities of about 3.1–3.3 km/s were found. The estimates for the average crustal Vs beneath Songpan-Ganzi and Qinling terranes is 3.5–3.6 km/s, which is in agreement with seismic refraction/reflection results (Liu et al., 2006). The crust below 20 km in the Songpan-Ganzi

and Qinling terranes has an average shear wave velocity of 3.6– 3.7 km/s and the high-velocity (Vs ≥ 4.0 km/s) lowermost crustal layer is either very thin or absent (≤4 km). These slow seismic velocities in the middle-to-lower crust were also reported by previous seismic refraction/reflection studies (Liu et al., 2006; Wang et al., 2003; Zhang et al., 2008), and coincide with observations from surface wave tomographic studies (Bao et al., 2013; Jiang et al., 2014; Li et al., 2014a; Yang et al., 2012) and joint inversion of receiver function and surface wave (Deng et al., 2015; Liu et al., 2014; Zheng et al., 2016). The remaining stations, mostly located in the Ordos Basin, display very slow shear wave velocities (≤3.3 km/s) at shallow depths, which likely represent the presence of sediments. Deep seismic sounding studies (Jia et al., 2014; Liu et al., 2006) also indicate the occurrence of ~5km-thick sedimentary deposits below Ordos Basin. Our velocity models indicate a 10–15-km-thick upper crust with S-wave velocities of 3.2–3.6 km/s overlying a 20–30-km-thick lower crust with shear wave velocities that increase from 3.6 km/s to 4.1 km/s just above the Moho. The average crustal Vs for the entire crust and below 20-km depth is about 3.7 km/s and 3.8 km/s, respectively. This finding is consistent with the seismic refraction/reflection result of Liu et al. (2006) and is 0.1–0.2 km/s faster than shear wave velocities in that depth range beneath the Songpan-Ganzi and West Qinling terranes. On average, a high-velocity layer 2–6 km thick is also observed at the lowermost part of the crust. While in the Qilian Terrane, two crustal types belonging to NE Tibet and Ordos crusts are observed.

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Fig. 4. Examples of the joint inversion for station 61063 (left) and 62335 (right). The locations of the two stations are shown in Fig. 1. (a) Observed (blue line) and predicted (red line) receiver functions. The receiver functions consist of four ray parameter bins (bracketed). The ray parameter is indicated to the left of the waveforms. Receiver functions with Gaussian factors at 1.0 (top) and 2.0 (bottom) are shown in each ray parameter bins. (b) Observed (solid dot with error bar) and predicted (black line) surface wave dispersions. (c) The shear wave velocity model from the joint inversion (black line) and the starting model (gray dashed line).

4. Discussion 4.1. Crustal composition

Fig. 5. 2-D shear wave velocity model across the seismic profile (see Fig. 1) formed by the interpolation of 1-D Vs models from the joint inversion. The black triangles represent the stations. The white circles are crustal thicknesses estimated from the H-κ stacking analysis. Elevation variations along the profile are also shown (upper panel).

The S-wave crustal velocities and bulk crustal Vp/Vs do not allow for a detailed definition of the crustal composition, but can provide constraints on crustal lithologies of NE Tibet and the Ordos Basin. Based on a comprehensive analysis of laboratory measurements, Christensen (1996) concluded that the continental crust with felsic, intermediate and mafic lithologies have low (≤1.76), medium (1.76 ≤ Vp/Vs b 1.81) and high (1.81 ≤ Vp/Vs b 1.86) Vp/Vs values, respectively. Within the variety of lower crustal rocks, felsic granulite have lower shear wave velocities (b3.8 km/s), mafic granulite have a shear wave velocity around 3.9 km/s, while garnet-bearing granulites have higher shear wave velocities (N 4.0 km/s) (e.g., Christensen, 1996; Rudnick & Gao, 2003). Thus, the low to moderate Vp/Vs values (1.67–1.79) observed below the profile are probably associated with average felsic to intermediate crust compositions, and the mafic layer in the lower crust is likely thin or absent (Li et al., 2006; Jiang et al., 2014; Vergne et al., 2002; Wang et al., 2016). The absence of high-velocity mafic lower crust is also


obvious from our shear wave velocity model (Fig. 5). For the southwestern part of Ordos Basin, the average Vs of ~3.8 km/s is found below 20km depth, indicating an intermediate-to-felsic lower crust. This conclusion is consistent with results from wide angle reflection/refraction studies that suggest that the lower crust of the Ordos Block is characterized by low P-wave velocities (6.4–6.8 km/s) and Poisson's ratio (0.25) (Jia et al., 2014; Liu et al., 2006). However, our observations do not support the suggestion by Yu et al. (2012) that a thick (20 km) mafic lower crust exists beneath the Ordos Basin. In the Songpan-Ganzi, West Qinling and Qilian terranes, however, the lower crustal composition is more felsic than the Ordos Block, as indicated by the average Vs (~ 3.5–3.6 km/s) below 20-km depth. The more felsic lithologies of these terranes are also reflected in the low Pwave velocities (6.3–6.6 km/s) from the seismic refraction/reflection results (Liu et al., 2006; Zhang et al., 2008). Alternatively, the lower shear wave velocities in the lower crust under the Songpan-Ganzi and West Qinling terranes can be caused by lower crustal partial melting, which has been previously proposed from an analysis of seismic refraction/reflection (Liu et al., 2006), surface wave (Bao et al., 2013; Jiang et al., 2014; Li et al., 2014a; Yang et al., 2012) and joint inversions of RF and surface waves (Deng et al., 2015; Liu et al., 2014; Zheng et al., 2016) data. However, no significant partial melting exists under NE Tibet (Li et al., 2006; Pan and Niu, 2011; Tian and Zhang, 2013; Vergne et al., 2002), which is characterized by low to medium crustal Vp/Vs values. 4.2. Implications for crustal thickening beneath NE Tibet The mechanisms responsible for Tibetan Plateau crustal thickening have been the subject of long lasting debates. Several distinct models have been invoked to explain it: (1) crustal channel flow (Clark and Royden, 2000); (2) vertically coherent thickening model (e.g., England and Houseman, 1986); and (3) upper crustal thickening model (e.g., Meyer et al., 1998). Clark and Royden (2000) pointed out the absence of upper crustal shortening across much of the eastern plateau margin and suggested that crustal deformation occurs mainly within the lower crust by ductile flow. The crustal LVZs observed by surface wave studies (Bao et al., 2013; Jiang et al., 2014; Li et al., 2014a; Yang et al., 2012) and joint inversions of RF and surface waves (Deng et al., 2015; Zheng et al., 2016) have also been interpreted as mid-to-lower crustal channel flow. However, a high crustal Vp/Vs is predicted (Li et al., 2006; Jiang et al., 2014; Wang et al., 2016) if the crustal shorting was a result of extrusion of the weak and ductile mid-to-lower crustal rocks from the central plateau into NE Tibet (Clark and Royden, 2000). This prediction is inconsistent with the low average crustal Vp/Vs values observed in this region. An alternative explanation of the thick crust is that the lithosphere is in the process of uniform thickening (e.g., England and Houseman, 1986). This vertically coherent deformation (England and Houseman, 1986) is shown by seismic anisotropy beneath NE Tibet, inferred from shear wave splitting (Chang et al., 2008; Eken et al., 2013) and receiver function studies (Wang et al., 2016) that show the fast directions are roughly in agreement with either the maximum extension or shear directions (Wang et al., 2008; Chang et al., 2017). The process of uniform thickening alone, however, fails to explain our observations that mafic layers in the lower crust are absent under NE Tibet. Our observations of low Vp/Vs and low Vs in the middle and lower crust below NE Tibet (Fig. 5 and Table 1) are consistent with the interpretation that the thick crust beneath NE Tibet was the consequence of significant thickening in the upper and middle crust by folding and thrusting (e.g., Meyer et al., 1998; Vergne et al., 2002). Another explanation for the lack of mafic lower crust was the removal of mafic crust by delamination (Ji et al., 2009; Liu et al., 2006; Vergne et al., 2002; Wang et al., 2014). This conclusion is consistent with results from Pn and Sn wave and surface wave tomographic studies that indicate the high-velocity upper mantle lid is absent beneath northern Tibet (Li et al., 2013a, 2013b; Pei et al., 2007).

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4.3. Low-velocity zone in the upper crust At many stations within the NE Tibetan Plateau, an obvious low-velocity (Vs b 3.3 km/s) layer in the upper crust was found (Fig. 5). These observations are consistent with seismic refraction/reflection results in NE Tibet by Wang et al. (2003, 2007) and Zhang et al. (2008), which revealed LVZs with Vp b 6.0 km/s in the upper crust above 20 km. Wang et al. (2007) suggested that these layers may be associated with active ductile detachment within relatively hot crust. Other seismic refraction/wide angle reflection profile (Liu et al., 2006) across these terranes also show obvious upper crustal LVZs beneath the Songpan-Ganzi and Qinling-Qilian terranes, but their depth (15–25 km) is slightly deeper than our results. It is possible that there are larger uncertainties regarding the depth of intracrustal discontinuities because the secondary reflection in the seismic refraction records is weak and not clear. This low-velocity layer in the upper crust has been found farther east and south and using the same method as this study (Deng et al., 2015; Liu et al., 2014). Those authors argued that the low-velocity layer they observed is a result of partial melting in the upper crust in the presence of water. It is not possible to draw an exact conclusion regarding the origin of the low-velocity layer from our velocity model; however, it does confirm the presence of a low-velocity layer in the upper crust. 4.4. Abnormal crustal thickness beneath the southwestern Ordos Basin Previous geophysical studies have shown that the Ordos Block is a stable tectonic block characterized by a thick and high-velocity lithosphere that extends to at least ~ 150-km depth (Bao et al., 2013; Li et al., 2013b). However, the southwest part of the Ordos Block has an inferred average crustal thickness of 46.7 ± 2.6 km (Fig. 5 and Table 1), which is ~ 4–5 km thicker than in the northern part (Jia et al., 2014; Li et al., 2014b; Liu et al., 2006; Teng et al., 2013; Wang et al., 2016) and similar to beneath the active NE Tibetan Plateau. It is reasonable to suggest that the thick crust root beneath the southwest Ordos Block is related to ongoing India–Asia collision. The shear wave splitting analysis of Yong and Chen (2016) revealed that large delay times with NW–SE fast polarization directions in NE Tibet are continuous across the boundary into the southwestern part of the Ordos Block, where the shear splitting parameters are also obviously different from those in the northern Ordos Block (Chang et al., 2017). Based on these observations, Yong and Chen (2016) suggested that the old cratonic lithosphere of the southwestern Ordos Block is currently being replaced with hot Tibetan asthenosphere at depth. This hypothesis is supported by recent surface wave tomography (Li et al., under review), where the maximum lithospheric lid velocities for the northern Ordos Block are slightly higher than in the southern part, suggesting a relatively weak strength in the southern Ordos Block. 5. Conclusions In this study, we generated new estimates of crustal thickness, mean crustal Vp/Vs values and crustal shear wave velocity distributions for 34 portable seismic stations using the H-κ stacking technology and joint analysis of receiver functions and Rayleigh-wave dispersion velocities. Furthermore, the 1-D shear wave velocity models were combined to produce 2-D images beneath a seismic profile, which revealed the complexities of the crustal structure beneath NE Tibet and the southwestern Ordos Block. Our results map the crustal thickness variations both for NE Tibet and the Ordos Block. The systematic spatial variations in crustal thickness are evidenced in previous receiver function and seismic refraction/reflection studies. This study confirmed a previously proposed low-velocity zone in the upper crust beneath the NE Tibetan Plateau and abnormal crustal thickness beneath the southwestern Ordos Basin. The low to moderate crustal Vp/Vs values are indicative of a felsic to intermediate crustal composition. However, there are no significant

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differences in the mean crustal Vp/Vs values for different terranes along this profile. This is different from what was observed by Wang et al. (2016) in this region, who argued that the average Vp/Vs values beneath the Ordos Basin are obviously larger than in NE Tibet. Northeastern Tibet is characterized by a predominantly felsic crust with a likely absent mafic lower crust, suggesting that crustal deformation in NE Tibet was the result of mafic lower crustal delamination or/ and felsic upper and middle crustal thickening rather than other processes, such as vertically coherent thickening or crust thickening by ductile flow. Low shear wave velocities are observed below 20-km depth beneath the Ordos Block, which is contrary to the interpretation of Yu et al. (2012) that thick mafic lower crust is still present beneath the Ordos Basin.

Acknowledgments We thank the editor of Tectonophysics and anonymous reviewers for insightful suggestions and comments that helped to improve this paper. This study was supported by the China National Special Fund for Earthquake Scientific Research in Public Interest (Grant No. 201308011) and NSFC (Grant Nos. 41474072, 41404069 and 41474089). Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.tecto.2017.03.020.

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