Terrestrial heat flow in the baiyinchagan sag, erlian Basin, northern China

Geothermics 86 (2020) 101799

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Terrestrial heat flow in the baiyinchagan sag, erlian Basin, northern China a

b,a,c,

Jiong Zhang , Shaopeng Huang Wei Xua

d

e

f

a

*, Yinhui Zuo , Yongshui Zhou , Zhi Liu , Wentao Duan ,

T

a

School of Human Settlements and Civil Engineering, Xi’an Jiaotong University, Xi’an, 710049, China Institute of Deep Earth Sciences and Green Energy, Shenzhen University, Shenzhen, 518060, China Department of Earth and Environmental Sciences, University of Michigan, Ann Arbor, MI, 48109-1005, USA d State Key Laboratory of Oil and Gas Geology and Exploitation, Chengdu University of Technology, Chengdu, 610059, China e Sinopec Zhongyuan Oil Field, Puyang, 457001, China f College of Geography and Environment, Baoji University of Arts and Sciences, Baoji, 721013, China b c

ARTICLE INFO

ABSTRACT

Keywords: Terrestrial heat flow Rock thermal conductivity Erlian Basin

The Baiyinchagan Sag of the Erlian Basin is a Mesozoic continental rift basin located in the eastern part of the Central Asian Orogenic Belt between the North China Craton and the Siberian Craton. The sedimentary fill comprises mainly the Cretaceous sandstones and mudstones underlying Cenozoic succession. Knowledge regarding the factors affecting the tectonic evolution of the Baiyinchagan Sag is lacking. The objective of this study is to obtain new terrestrial heat flow measurements in the Baiyinchagan Sag for constraints on its thermotectonic history. We obtained continuous temperature logs from three boreholes, namely Xi24, Xi3-76 and Xi30 with depths of 1230, 1660 and 2160 m, respectively. Additionally, we gauged the thermal conductivity of the lithologic samples of the main sedimentary rocks. Our analysis indicates that subsurface temperatures are affected by the convection of groundwater and subsurface warming. After removing the influence of groundwater and climatic effects, the subsurface temperature gradients of the Cretaceous Duhongmu and Tenggeer Formations range from 40.4 to 42.1 ℃ km−1. We performed water saturation correction for the measured thermal conductivity data. The thermal conductivities of the Saihantala, the Duhongmu, the Tenggeer, and the Aershan Formations were found to be 1.15, 2.13, 1.99, and 2.15 W m-1 K−1, respectively. Vertical Peclet number analysis of the temperature-depth profile of Xi3-76 borehole with a Peclet number of 0.16 indicates that the conductive heat transfer dominates in the component of the vertical heat transfer. The three new heat flow values range from 80.8 to 89.7 mW m-2, with a mean of 84.9 ± 3.4 mW m-2, significantly higher than the global and regional average whereas consistent with its deep thermotectonic regime.

1. Introduction The Baiyinchagan Sag (BS) (Fig. 1) (longitude E 107°30′ to E 109°10′; latitude N 41°50′ to N 42°30′) is situated in the west of the Erlian Basin (EB) in Inner Mongolia, with an area of 3200 km2. It is located in the eastern part of the Central Asian Orogenic Belt (CAOB) between the North China Craton (NCC) and the Siberian Craton. Petrochemical evidence (Yang et al., 2017) shows that the basement of the Baiyinchagan Sag belongs to the northern margin of the NCC, a destroyed cratonic margin during the Paleozoic and Mesozoic. An earlier report based on oil-field formation testing data estimated the geothermal gradient of the Erlian Basin to be in the range of 31 to 43 ℃ km−1 (Ren et al., 2000). The oil-field production data derived heat flow values range from 43.1 to 66 mW m-2 with a mean of 55.5 mW

⁎

m-2 (Xiao et al., 2004), which is lower than the average value of 61.5 ± 13.9 mW m-2 for the continental China (Jiang et al., 2019). However, recent studies have shown that the Erlian Basin is of a significantly higher heat flow. Zuo et al. (2016), based on oil field production data too, reported a heat flow of 75.5 mW m-2 for the Baiyinchagan Sag, whereas Xu et al. (2018) reported that the heat flow of the Uliastai Sag in the northeast of the Erlian Basin is 86.3 ± 2.3 mW m-2. These newly published results show that the Erlian Basin may not be a low or medium heat flow region, but a high heat flow and active tectonic region. The tectonic activities of the Erlian Basin in the Mesozoic and Cenozoic must have greatly influenced the dynamic evolution of the NCC. The transformation of the tectonic regime of the NCC began from 150-140 Ma to approximately 110-100 Ma, peaking at ∼120 Ma (Zhai

Corresponding author at: Institute of Deep Earth Sciences and Green Energy, Shenzhen University, Shenzhen, 518060, China. E-mail address: [email protected] (S. Huang).

https://doi.org/10.1016/j.geothermics.2019.101799 Received 25 April 2019; Received in revised form 30 September 2019; Accepted 29 December 2019 0375-6505/ © 2020 Elsevier Ltd. All rights reserved.

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Fig. 1. Tectonic setting of the study area (modified from Huang and Zhao (2006); Santosh (2010) and Xia et al.(2017)). CAOB- Central Asian Orogenic Belt; NCCNorth China Craton; EB- Erlian Basin; YB- Yinshan Block; IMSZ- Inner Mongolia Suture Zone; OB- Ordos Block; Yellow line- the line of 108 °E longitude (Used in the discussion section); CJ- Chuanjing Depression; WLCB- Wulanchabu Depression; MNT- Manite Depression; SNT- Sunite Uplift; TGE- Tenggeer Depression; WNTWunite Depression (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article).

et al., 2004). The destruction of the NCC peaked at about 125 Ma (Zhu et al., 2011, 2012). The Pacific plate retreat and back-arc extension started around 110 Ma, following magmatism reactivation (Sun et al., 2008). In comparison, the Erlian Basin began its pre-rift stage at 156 Ma (Bonnetti et al., 2014), with multistage tectonic-thermal events, strong magmatic activities, and structural deformation that occurred in the study area under lithospheric alteration from 125 to 107 Ma. In particular, a strong magmatic event occurred at around 110 Ma. The timeline of the tectonic events of the Erlian Basin is consistent with the destruction of the NCC. Previous tomographic studies have indicated

that the lithosphere thickness of the present northern margin of the NCC is between 80 and 100 km (Huang et al., 2009). P wave receiver functions reveal that the Moho depth is about 44 km beneath the northern margin of the NCC (Tian et al., 2011). The thickness of the lithospheric mantle ranges from 36 to 56 km, which is more strongly modified and thinner than a stable cratonic block. However, what occurred during the lithosphere coupling process is unclear, and the results of recent studies do not agree well with each other. Therefore, this problem needs to be further elucidated. Heat flow measurements can help constrain the calculation of the subsurface temperature 2

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distribution and lithospheric thermal structure, which are crucial for the reconstruction of a past thermotectonic history. The objective of this study is to provide a more accurate heat flow for the Baiyinchagan Sag in the western Erlian Basin. The new heat flow values are based on systematic borehole temperature and thermal conductivity measurements. In particular, the heat flow calculation in our study considers the influence of the groundwater advection and transient effect of ground surface temperature variation. In this paper, section 2 describes the geological background of the Baiyinchagan Sag. Section 3 is dedicated to the description of the borehole temperature measurements in the Baiyinchagan Sag. Section 4 focuses on rock thermal parameters. Section 5 presents the calculation of the eventual heat flow for the Baiyinchagan Sag. The discussion and conclusions are presented in section 6.

samples from the temperature-logged boreholes, thermophysical properties of rock samples from nearby boreholes of Weng1, Da6, Xi2, and Wu1 were measured to represent the rock thermophysical properties of the whole Baiyinchagan Sag. The locations of the boreholes are shown in Fig. 2a. Borehole temperatures were measured using a Matrix Logging System, made by Mount Sopris Instrument Co., Inc. The system consists a 42.9 mm diameter sensor (maximum temperature bound: 176 ℃, maximum pressure bound: 124 MPa) and a 5000 m long cable (2.54 mm diameter single conductor). It allows temperature recording to a sensitivity of 0.01 ℃ at an accuracy of 0.1 ℃. The temperatures were recorded at depth intervals of 0.1 m, and the downward speed was 9 m min−1 to ensure sufficient time for the sensor to record the temperature at different depths. Details regarding these boreholes are presented as follows: (1) Drilling for Xi24 began on April 10, 2010, and completed on April 28, 2010. Its depth is 2080 m and diameter is 215 mm. The bottom of Xi24 belongs to the Paleozoic Formation. (2) Drilling for Xi30 began on March 30, 2010, and completed on April 29, 2010. Its depth is 2420 m and the diameter is 215 mm. The bottom of Xi30 belongs to the Aershan Formation. (3) Drilling for Xi3-76 began on May 9, 2006, and completed on July 10, 2006. Its depth is 2190 m and diameter is 215 mm. The bottom of Xi3-76 belongs to the Aershan Formation. However, due to some obstruction inside the boreholes, our temperature logging instrument was unable to reach the bottom of these three boreholes. The temperature-depth profiles of these three boreholes are displayed in Fig. 3. At the shallow depths smaller than 200 m, subsurface temperatures are usually influenced by climate change and near-surface environmental factors including ground water flow. At greater depths, the variation of temperature tends to be mainly constrained by heat conduction and is almost linear. The temperature-depth profiles of Xi30 and Xi3-76 boreholes are unrepresentative of formation temperature down to about 200 m depths because they were measured in the air column above the water level. Excluding the upper 200 m sections, the depth and corresponding temperature ranges are 200−2160 m and 13.8–89.0 ℃ for Xi30, 200−1230 m and 14.6–59.9 ℃ for Xi24, and 200−1660 m and 13.1–65.4 ℃ for Xi3-76, respectively. The temperatures at 200 m of these three boreholes range from 13.1 to 14.6 ℃, suggesting that the geothermal environments of these three boreholes at 200 m are basically similar. However, the upper 500 m section of the borehole temperature of Xi24, located in the Baiyinwengte Structural Belt, is shown to be affected by some near surface processes. The temperature-depth profile of Xi30 is consistent with that of Xi3-76 at the range 200−1300 m. The temperatures at 1300 m depth of both boreholes is 55 ℃. These two boreholes are located in the Western SubSag.

2. Geological background The Erlian Basin is located in the central-eastern Inner Mongolia and consists of five depressions (the Chuanjing Depression, the Wulanchabu Depression, the Manite Depression, the Tenggeer Depression, and the Wunite Depression) and an uplift (the Sunite Uplift) (Fig. 1). The Baiyinchagan Sag is a secondary structural unit of the Chuanjing Depression and is amongst the areas with high oil and gas resources potential in the Erlian Basin. Since the Paleozoic Era, the Erlian Basin had been affected by the closure of the Paleo-Asian Ocean (Safonova et al., 2011), and has gone through plate consolidation and orogeny. From the late Mesozoic, along with the decline of the Paleo-Asian Ocean tectonic domain and the closure of the Mongolian-Okhotsk Ocean (Tang et al., 2016; Wilde and Zhou, 2015), the subsidence and sedimentation of the Erlian Basin has been mainly attributed to the back-arc extension associated with the Paleo-Pacific Plate subduction (Xiao et al., 2001). From north to south, the main structural units of the Baiyinchagan Sag are the Tala Fault Belt, the Western Sub-Sag, the Baiyinwengte Structural Belt and the South Slope (Fig. 2a). Fig. 2b shows the stratigraphic structure of the Baiyinchagan Sag. The basement of the Baiyinchagan Sag consists the Paleozoic strata (Pz), and had undergone substantial erosion; the sedimentary environment was dominated by neritic facies and marine/continental facies (Cui et al., 2011). Triassic and Jurassic strata are absent from the Baiyinchagan Sag. The sedimentary fill comprises the Aershan (K1ba, include K1ba1 and K1ba2), the Tenggeer (K1bt), the Duhongmu (K1bd, include K1bd1, K1bd2, and K1bd3), the Saihantala (K1bs), the Erliandabusu (K2er) Formations, and the overlying Cenozoic succession. It has experienced three phases of tectonism since the early Cretaceous (Bonnetti et al., 2014) including (1) (135-103 Ma) a rift stage (deposition of the Aershan Formation to the Duhongmu-1 Formation when the sedimentary environment was dominated by fluvial, shore facies, and Lacustrine and characterized by intense faulting); (2) (103-95 Ma) a transition stage from synrift to postrift (deposition of the Duhongmu-2 Formation to the Saihantala Formation when the sedimentary environment was dominated by lacustrine and characterized by weak faulting and a wide lake basin); and (3) (95 Ma-present) postrift and uplifting stages when the strata were thinned in the fluvial environment (Zuo et al., 2016).

3.2. Geothermal gradient Geothermal gradient is the rate of underground temperature variation with respect to the depth change. In this paper, we first calculated a borehole temperature gradient using the least squares linear regression method (Powell et al., 1988) for the whole temperaturedepth profile excluding the upper 200 m. We further calculated temperature gradients at a depth interval of 20 m and subsequently mean formational temperature gradients (Table 1). Utilizing the temperature gradients at depth intervals of 20 m, we calculated the observed heat flow density versus depth in section 6.1. For the borehole Xi24, the average temperature gradient was determined to be 45.4 ℃ km−1 with a correlation coefficient of 0.994 and an intercept of 5.87 ℃. The temperature-gradient values ranged from 28.8 to 79.6 ℃ km−1. In the Saihantala Formation (200−360 m), the temperature-gradient values ranged from 32.3 to 79.6 ℃ km−1, and the mean ± SD was 41.1 ± 10.0 ℃ km-1. In the Duhongmu Formation (360−1230 m), the temperature-gradient values ranged from 28.8 to 59.5 ℃ km−1, and the mean ± SD was 45.3 ± 11.3 ℃ km-1.

3. Borehole temperature measurements 3.1. Temperature as a function of depth Temperatures of three deep boreholes (Xi24, Xi3-76 and Xi30 with depths of 1230, 1660, and 2160 m, respectively) in the Baiyinchagan Sag were measured by continuously logging. To allow sufficient time for the borehole temperature to reach a steady state, all of the selected boreholes were kept away from perturbation for more than one year. Xi24 is located in the Baiyinwengte Structural Belt, the other two boreholes are located in the Western Sub-Sag. Due to the lack of core 3

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Fig. 2. Borehole location (a) and cross section A-A’ (b) of the Baiyinchagan Sag (Zuo et al., 2016).

For the borehole Xi30, the average temperature gradient, the correlation coefficient, and the intercept are 39.3 ℃ km−1, 0.998, and 4.01 ℃, respectively. The temperature-gradient values ranged from 13.2 to 68.7 ℃ km−1. In the Saihantala Formation (200−400 m), the temperature-gradient values ranged from 13.2 to 68.4 ℃ km−1, and the mean ± SD was 31.8 ± 15.8 ℃ km-1, whereas in the Duhongmu Formation (400−1430 m), the temperature-gradient values ranged from 16.6 to 51.4 ℃ km−1, and the mean ± SD was 39.3 ± 13.9 ℃ km-1. In the Tenggeer Formation (1430−2160 m), the temperaturegradient values ranged from 34.0 to 51.9 ℃ km−1, the mean ± SD was 40.9 ± 3.6 ℃ km-1. For the borehole Xi3-76, the average temperature gradient, the correlation coefficient, and the intercept are 36.6 ℃ km−1, 0.989, and 5.89 ℃, respectively. The temperature-gradient values ranged from 22.5 to 74.5 ℃ km−1. In the Saihantala Formation (200−440 m), the temperature-gradient values ranged from 30.7 to 74.5 ℃ km−1, and the mean ± SD was 33.1 ± 3.5 ℃ km-1. In the Duhongmu Formation (440−1230 m), the temperature-gradient values ranged from 25.1 to

Fig. 3. Temperature profiles versus depth.

Table 1 Linear regression analysis of the geothermal profiles and the range of temperature gradients at depth intervals of 20 m in different formations. Borehole

Linear least squares regression method Slope (℃ km−1)

Intercept

Linear correlation coefficient

Xi24

45.4

5.87

0.994

Xi30

39.3

4.01

0.998

Xi3-76

36.6

5.89

0.989

Strata (m) K1bs (200–360) K1db (360–1230) K1bs (200–400) K1db (400–1430) K1bt (1430–2160) K1bs (200–440) K1db (440–1230) K1bt (1230–1660)

4

Temperature gradient for depth intervals of 20 m (℃ km−1) Range

Mean ± SD

32.3–62.6 28.8–75.6 13.2–68.4 16.6–68.7 34.0–51.9 22.5–74.5 33.1 – 3.5 22.5–51.6

41.1 45.3 31.8 39.3 40.9 33.1 38.9 33.6

± ± ± ± ± ± ± ±

10.0 11.3 15.8 13.9 3.6 3.5 13.9 8.8

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Fig. 4. Schematic diagrams of the subsurface thermal regime under the condition of groundwater advection (a) and climate change (b).

36.3 ℃ km−1, and the mean ± SD was 38.9 ± 13.9 ℃ km-1. In the Tenggeer Formation (1230−1660 m), the temperature-gradient values ranged from 22.5 to 51.6 ℃ km−1, and the mean ± SD was 33.6 ± 8.8 ℃ km-1. In general, the temperature-gradient values are at the shallower depth lower above the Saihantala Formation and the Duhongmu Formation. Such variations were not obvious at the lithologic interfaces between the Duhongmu Formation and the Tenggeer Formation. This difference is possibly due to the groundwater convection in the Baiyinchagan Sag caused by the oil production activities from the surrounding boreholes. In particular, the temperature gradient of Xi376 shows a decline at 840 m at the Duhongmu Formation. Therefore, the effect of groundwater advection needs to be evaluated for heat flow calculation.

km−1; U = vc0 0 / c ,(negative U denotes upward groundwater flow), where v is the vertical groundwater flow in m a-1, c0 0 is the heat capacity of the water and c is the heat capacity of the aquifer; b is the rate of change in surface temperature in ℃ year−1; erfc is a complementary error function; α is the thermal diffusivity, which is 6.5 × 10-7 m2 s-1. The modeling is limited for semi-infinite layers with only vertical conduction and convection, and vertical groundwater flow is assumed to be constant with depth. The details about the relationship among t, U, b and z have been presented by Taniguchi (Taniguchi et al., 1999). Utilizing the 300−500 m section temperature-depth profile of Xi30, we estimated the steady state geothermal gradient G to be 41.5 ℃ km−1, which is the slope of the borehole T-z linear regression (Fig. 5). Via trial and error by adjusting the parameters t, U, b and T0, we tried to match the profile observed. The calculated profile with G = 41.5 ℃ km−1, t = 200 a, U = 0.9 m a-1, b = 0.04 ℃ year-1 and T0 = 10 ℃ matches the profile observed at the Xi30 borehole reasonably well. The ground surface warming rate b = 0.04 ℃ year−1, the warming reported from many locations over the world (Taniguchi et al., 1999; Barkaoui et al., 2013; Bayer et al., 2016; Westaway and Younger, 2016). Taniguchi et al. (1999) reported an U of 0.4-0.6 m a-1 for the Musashino terrace west to the Tokyo Bay, which is of the same order as the result of our study. Since the borehole temperatures are affected by groundwater flow and climate change in this region, calculating the subsurface temperature gradient by using the whole temperature-depth profiles wouldn’t be appropriate. To derive the heat flow, we choose the temperaturedepth profiles below 200 m to calculate the temperature gradients, to avoid the influence of climate change. We exclude the segments that are affected by groundwater and calculate the subsurface temperature gradient for different formations. Additionally, we correct the temperature-depth data and calculate the heat flow of Xi3-76 in Section 5.2.

3.3. Influence of climate change and groundwater advection Two most commonly seen perturbations to the underground temperature distribution are groundwater advection (Uchida et al., 2003) and climate change (Huang et al., 2000). Fig. 4 shows the schematic diagrams of the patterns of these two effects. In order to calculate a more reliable heat flow, it is essential to eliminate perturbed temperature-depth sections if avoidable, or make appropriate correction if unavoidable. Temperature-depth profiles are mostly affected by climate change at shallow depths. In this paper, only the temperature-depth profile of Xi30 borehole (Fig. 3) has a typical U-shape in the upper 200 m depth apparently due to the climatic effect, whereas the water levels of Xi24 and Xi3-76 were at about 200 m depth making the temperature measurements from the Xi24 and Xi3-76 unsuitable for analysis of climatic effect (Pollack and Huang, 2000). In order to evaluate the dual effect of climate change and groundwater flow on the subsurface temperature at the shallower depths, we analyzed the upper 500 m temperature profile of Xi30 with the onedimensional heat conduction-advection equation of Carslaw and Jaeger (1959)

T (z, t ) = T0 + G (z

4. Thermal conductivities of core samples

Ut ) + {(b + GU )/2U }

Core sampling is performed only for a few proposes by the geological exploration department of the Zhongyuan Oilfield because it is too expensive. Unfortunately, no coring could be performed in the boreholes from which we have systematically measured temperatures. Therefore, the rock thermal conductivities were measured on the cores of four boreholes of Weng1, Wu1, Da6, and Xi2 in the Baiyinchagan Sag (Fig. 2b), which were previously acquired by the Zhongyuan Oilfield.

× [(z + Ut ) eUz / erfc {(z + Ut )/2( t )1/2} + (Ut

z ) erfc {(z

Ut )/2( t )1/2}]

(1)

where T(z,t) in ℃ is the temperature at given depth z in m and given time t in year since the onset of surface temperature change; T0 is surface temperature in℃; G is the steady state geothermal gradient in ℃ 5

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Fig. 5. Linear regression analysis of the temperature-depth profiles of Xi30 borehole at upper 300−500 m depth (a). Comparison of calculated temperature-depth profiles with groundwater flow and climate change, and profile observed of Xi30 borehole at upper 500 m depth (b).

and 2.19 W m-1 K-1 for the Saihantala, the Duhongmu, the Tenggeer and the Aershan Formations, respectively. Thermal conductivity of the rocks depends on many factors, such as pressure, temperature, porosity, and saturated fluids. Our results show that rock thermal conductivity decreases with increase of temperature and is positively correlated with pressure; the sum of the effects of temperature and pressure may be very small (Abdulagatova et al., 2009). Due to the porous nature of underground rocks, the thermal conductivities of rocks under dry and water-saturated conditions differ significantly. Thus, it is essential to perform water saturation correction. Abdulagatova et al. (2009) summarized many ways of water saturation correction including geometric mean and arithmetic mean. The porosity of sandstone samples ranges from 5 % to 27 % in the Baiyinchagan Sag, and that of the mudstone samples is lower than 1 % and thus is ignored. Guo et al. (2017) found that the deviation in thermal conductivity is up to 20 % if the porosity is higher than 10 %. Here, we used the arithmetic mean method for water saturation correction. The corrected thermal conductivity λ is represented as

Our measurement principle is that core samples should be representative of every major formation presented in the basin, including mudstone, sandstone, schist and basalt. In this paper, the thermal conductivity was tested at a constant temperature of about 24 ℃. We measured the thermal conductivities of 41 core samples, of which detailed information is shown in Table 2. The instrument used is a high precision noncontact optical scanning system (Popov et al., 1999a, b). Its accuracy of measurement is ± 3 % for the range from 0.2 to 25 W m−1 K-1. Measurements were performed on dry cores under ambient P-T conditions. The thermal conductivity of the standard sample is 2.37 W m−1 K-1. The measured thermal conductivities of the 41 samples are in the range of 0.89–4.91 W m−1 K-1. The measurements are 1.10–2.70 W m−1 K-1 with a mean of 1.86 ± 0.45 W m−1 K-1 for the 18 mudstone samples, 0.89–3.14 W m−1 K-1 with a mean of 1.81 ± 0.54 W m−1 K-1 for the 18 sandstone samples, 3.11–4.91 W m−1 K-1 with a mean of 4.07 ± 0.78 W m−1 K-1 for the 3 schist samples, and 1.23-1.40 W m−1 K-1 with a mean 1.29 ± 0.09 W m−1 K-1 for the 2 basalt samples. The rock thermal conductivity measurements show that the average thermal conductivity of schist is the largest and that of basalt is the smallest, whereas the average thermal conductivities of mudstone and sandstone are almost the same. The average thermal conductivities of mudstone are 1.17, 1.89, 1.88 and 2.04 W m-1 K-1 for the Saihantala, the Duhongmu, the Tenggeer and the Aershan Formations, respectively. The average thermal conductivities of sandstone are 1.11, 2.18, 1.94

= m0

m

=

× (1 m

× (1

)+

w

)+

(2)

× a

(3)

×

where λm is the thermal conductivity of the matrix; λw is the thermal conductivity of pore fluid; λ m0 is the observed thermal conductivity; λa is the thermal conductivity of air; and φ is the porosity in %. The

Table 2 The thermal conductivity observed, calibration for water-saturation and weight mean values for different lithologies. Strata

Shale percentage content (%)

K1bs

67.2

K1bd

53.3

K1bt

62.8

K1ba

45.2

Lithology (n)

ф

λm0 (W m−1 K-1)

λsat (W m−1 K-1)

Average (W m−1 K-1)

Sandstone (3) Mudstone (3) Sandstone (3) Mudstone (5) Sandstone (11) Mudstone (4) Sandstone (1) Mudstone (6)

0.06-0.14 0 0.08-0.27 0 0.06-0.15 0 0.1 0

0.89–1.50 1.1–1.24 1.49–3.14 1.51–2.20 1.72–2.28 1.45–2.79 2.19 1.64–2.69

0.92–1.57 – 1.56–3.26 – 1.75–2.36 – 2.24 –

1.11 1.17 2.21 1.89 2.07 1.88 2.24 2.04

The symbol “-” means no calibration. 6

Weighted mean (W m−1 K-1) 1.15 2.13 1.99 2.15

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thermal conductivity of the pore fluid is 0.6 W m−1 K-1, and the thermal conductivity of air is 0.03 W m−1 K-1. According to the percentage of sandstone and mudstone in the same stratum, the thermal conductivity of a stratum is calculated by equation

=

s

× Ps +

n

subsurface warming and downward flow in shallow formation. Compared with the calculated values, the means of three boreholes are significantly lower in the Saihantala Formation. The heat flow density versus the Duhongmu Formation of three boreholes show a decline trend from upper to under segment. We conclude that the groundwater advection is significantly in the Duhongmu Formation.

(4)

× Pn

where λs is the thermal conductivity of sandstone; λn is the thermal conductivity of mudstone; Ps is the percentage content of sandstone; and Pn is the percentage content of mudstone. The data regarding porosity and amount percentage were provided by Zhongyuan Oil Field Geology Institute. A comparison of the observed and calibrated thermal conductivities is shown in Table 2. According to the different lithology ratios (Table 2), the rock thermal conductivities of the Saihantala, the Duhongmu, the Tenggeer and the Aershan Formations determined by Eq. (4) are 1.15, 2.13, 1.99 and 2.15 W m−1 K−1, respectively. In addition, we use the same method to derive the thermal conductivities of the Saihantala, the Duhongmu, the Tenggeer and the Aershan Formations, which are 1.14, 2.03, 1.90 and 2.12 W m−1 K−1, respectively.

5.2. Groundwater effect Groundwater advection plays an important role in the transportation of geothermal heat. For a deep circulating system, groundwater activity includes vertical convection (upward or downward flow) and horizontal radial flow. The horizontal radial flow has less perturbation, while vertical convection affects the subsurface temperature distribution substantially. Bredehoeft and Papaopulos (1965) proposed the Peclet number (Pe) analysis method. Pe is defined as the ratio of heat advected and the heat conducted:

Pe =

5. Heat flow

Terrestrial heat flow is a crucial indicator of the Earth’s interior thermal state. Terrestrial heat flow density, or heat flow in short, represents the heat transmitted from inside of the Earth to the per unit area of the surface and reflects the energy balance between various geodynamic processes in the Earth’s interior. Heat flow is usually determined by using

Q=

=

( c )f v T

( ) T L

=

Qadvective Qconductive

(6)

where ( c )f is volumetric specific heat capacity of water, 4.1665 MJ m−3 K-1; v is Darcy velocity (negative downward) in mm a-1; and L = z b z a is the depth interval of the vertical groundwater flow, where z a and z b are the depths of the roof and floor of the groundwater flow permeable layer, respectively. Alternatively, the ratio Pe can be L determined from a linear regression of temperature gradient versus temperature (Mansure and Marshall, 1979). Clauser and Villinger (1990) proposed a logarithm linear regression method to determine the Pe ratio:

5.1. Calculation of heat flow

dT dz

( c )f vL

L

Q (z ) = Q (z a ) e

(5)

where Q is the heat flow in mW m−2, and dT is the geothermal gradient dz in ℃ km-1. The negative symbol indicates that the direction of heat flow is opposite to that of the geothermal gradient. Firstly, for the temperature-depth profiles of Xi24, Xi30 and Xi3-76, groundwater convection effects were removed. Next, the temperaturedepth profiles of the Duhongmu Formation and the Tenggeer Formation were extracted, which were mainly dominated by heat conduction. Heat flow was calculated by multiplying the least-square temperature gradient with thermal conductivity (Table 3). The results showed that the range of geothermal gradients was 40.4–42.1 ℃ km−1. The calculated heat flow values range from 80.9 to 89.7 mW m-2, with a mean of 84.9 ± 3.4 mW m-2. In our study, utilizing temperature gradients at depth intervals of 20 m and the measured conductivities of the core samples from different lithology, we plot the measured heat flow density versus depth profiles for the three boreholes (Fig. 6). The main lithology of the Saihantala Formation is sandstone-mudstone interbed. The mean measured conductive components of the Saihantala Formation heat flow of Xi24, Xi30 and Xi3-76 are 46.9 ± 12.1, 36.3 ± 18.9 and 37.7 ± 4.2 mW m−2, respectively, which are affected by the

(z z a) × Pe L

ln (Q (z )) = ln (Q (z a ))

(7)

(z

z a) ×

Pe L

(8)

where Q (z a ) is the heat flow at z a ; Pe and Q (z a ) can be determined by L the linear regression of the logarithm of the heat flow density versus depth. Therefore, the theoretical temperature distribution in a homogeneous half-space can be derived from Eq. (9) (Mansure and Marshall, 1979) based on the values for z and Pe . L

T (z ) =

(

Q (z a )

( ) Pe L

(z

)×

exp

z a) ×

(z

z a) ×

Pe L

1 + T (z a ) × exp

Pe L

(9)

where T (z a ) is the actual temperature at z a . The distribution of the subsurface temperature of Xi3-76 is substantially deviated from linear distribution over the depth interval of Duhongmu Formation, appearing to be affected by groundwater. The abrupt change in the temperature-depth profile at the depth of 840 m (Fig. 3) suggests that there is an interface of two different groundwater flow systems (Fig. 7a). The system above the interface is much complex

Table 3 The temperature gradient and heat flow of the three boreholes. Borehole

Longitude (E)

Latitude (N)

Strata

Depth interval (m)

Temperature gradient (℃ km−1)

Coefficient of determination

Thermal Conductivity (W m−1 K-1)

Heat flow (mW m−2)

Mean ± SD (mW m−2)

Xi24

107°50′29.8″

42°00′33.7″

Xi30

107°47′10.7″

42°01′46.5″

107°45′45.7″

42°02′22.1″

600-1230 400-700 1430-2160 440-700 1500-1660

40.4 40.7 40.9 42.1 40.6

0.998 0.993 0.999 0.997 0.999

2.13 2.13 1.99 2.13 1.99

86.1 86.7 81.4 89.7 80.8

84.9 ± 3.4

Xi3-76

K1bd K1bd K1bt K1bd K1bt

7

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Fig. 6. Lithology, 20 m interval temperature gradients, thermal conductivities and measured heat flow density versus depth of Xi24 (a), Xi30 (b) and Xi3-76 (c).

because the lithology is interbeded sandstone-mudstone, which is difficult to determine the roof and floor of the aquifer exactly. In contrast, the stratum between 860 m and 1260 m is sandstone, which can be treated as a groundwater flow permeable layer (Fig. 6c). Utilizing Eq. (5), the geothermal gradient dT/dz (20 m interval) and the thermal conductivity of the Duhongmu Formation, we could acquire the heat flow Q (z ) and the relationship between z and ln (Q (z )) . The results (Fig. 7a) show that for the water flow system of 860−1260 m depth range Pe = 0.4 × 10−3 m-1, Pe = 0.16, and Q (z a ) = 99.5 mW m-2, L which is the roof heat flow affected by groundwater advection and accordant with the measured heat flow of Xi3-76 Duhongmu

Formation. The calculated and observed temperatures for the depth interval 860−1260 m are shown in Fig. 7b. The average vertical Darcy velocity is v = 6.6 mm a-1. According to Eq. (7), we calculate the heat flow Q (z b) = 84.8 mW m-2 at the depth z b , which is the floor heat flow. 6. Discussion 6.1. Reliability analysis of high heat flow Early estimates based on oil field production data suggested the Erlian Basin is a cold basin with a heat flow substantially lower the

Fig. 7. Peclet number (Pe) analysis of the Xi3-76 Duhongmu Formation geothermal data. (a): ln (Q (z )) versus z profile showing two separated groundwater circulations systems separated at the depth around 850 m. (b): Observed and calculated temperatures for the depth interval of 860−1260 m. 8

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Fig. 8. Schematic illustration of geodynamics of the study region. See caption of Fig. 1 for the abbreviations of tectonic units and the profile location. (a) The resistivity section along 108 °E at a depth of 150 km (modified from Dong et al. (2014)). (b) P-wave velocity tomography along 108 °E at a depth of 600 km (modified from Santosh (2010); Santosh et al. (2010) and Tian et al. (2009)).

mean of the continental China. Our study shows the opposite. In this study, the effect of groundwater flow on the heat flow measurements is eliminated. Heat flow is calculated by utilizing the subsurface temperature-depth profile that is dominated by heat conduction. The range of the geothermal gradients is 40.4–42.1 ℃ km−1 as opposed to the 30.0–38.0 ℃ km−1 estimated by Xiao et al. (2004) using oil-field formation testing data. Obviously, there are differences between the systematically measured borehole temperatures and the formation testing temperatures. The temperatures measured from long-resting boreholes are more reliable for calculating the subsurface geothermal gradient. We measured the rock thermal conductivity of 41 core samples from the major strata of the Baiyinchagan Sag. Additionally, we considered the effects of groundwater saturation on thermal conductivity of porous rocks, which was substantial. The average value of the heat flow in the Baiyinchagan Sag is 84.9 ± 3.4 mW m-2, which is higher than the estimated heat flow of 75.5 mW m-2 based on formation-testing temperature data and the thermal conductivities from the adjacent Chagan Sag (Zuo et al., 2016). It is much higher than the heat flow 64.7 ± 8.9 mW m-2 of the Ordos Basin (Gao et al., 2018) and the global continental average heat flow (65 mW m-2) (Pollack et al., 1993). However, our result is consistent with the heat flow of 86.3 ± 2.3 mW m-2 recently reported for the Uliastai Sag in the northeast of the Erlian Basin (Xu et al., 2018). According to paleogeothermal research, the Baiyinchagan Sag experienced a much higher geothermal gradient in the Cretaceous period than present. Zuo et al. (2016) used apatite fission track and vitrinite

reflectance from oil wells to reconstruct the Mesozoic thermal history, and showed that the Baiyinchagan Sag might have had experienced geothermal gradients as high as 49.9–56.4 ℃ km−1 at the end of the deposition of the Saihantala Formation (100–95 Ma). At around 100 Ma, an abrupt heat flow increase in the Ordos Basin was revealed by the analysis of vitrinite reflectance of sedimentation from oil wells (Jiao et al., 2013). The timing of heat flow peak in the Ordos Basin is consistent with the time of geothermal gradient peak in the Baiyinchagan Sag. Our results indicate that the Erlian Basin is not a middle-low heat flow area as thought previously. The new measurements set a new geothermal constraint for the study of the tectonic evolution of the Erlian Basin. 6.2. Evidence of asthenosphere upwelling The Erlian Basin and the adjacent areas have recently undergone several geophysical investigations. P wave receiver functions reveal that the Moho depth is 44 km beneath the study area and 47−48 km beneath the Yinshan Block. The velocity of the P wave is 6.2 km s−1, and the thickness of the low velocity anomalous body is about 6−8 km within the crust around the study area (Teng et al., 2010; Tian et al., 2011). Recent petrochemical study on the Mesozoic bimodal volcanic rocks from the Yinshan Block (Guo et al., 2018) shows that the magmas parental to the basaltic rocks in this region were originated from partial melting of ancient metasomatized sub-continental lithospheric mantle, 9

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whereas the magmas parental to the rhyolites were formed by mixing of mantle-derived alkaline mafic melts and lower crust anatexis derived felsic melts. The result of Magnetotelluric sounding (Dong et al., 2014) further suggests a low resistance anomalous body (< 10 Ω m) in the crust beneath the Yinshan Block. The Curie Point isotherm lies at the depths of about 19−24 km in the study area (Xiong et al., 2016), and indicates that there may be magmatic activity. The crustal density structure of the gravity data showed that a heterogeneous low-density anomalous body existed in the crust (Wang et al., 2010). The lowdensity anomalous body is likely to be a mafic or ultramafic complex from the deep asthenosphere (Zhang et al., 2014). The high heat flow anomaly as revealed by our new measurements could have been caused by such a low-density body.

Declaration of Competing Interest The authors declared that they have no conflicts of interest to this work. We declare that we do not have any commercial or associative interest that represents a conflict of interest in connection with the work submitted. Acknowledgements This study was financially supported by the National Natural Science Foundation of China (41374089). We thank Shengbiao Hu, Yi Wang and Di Hu of the University of Chinese Academy of Sciences for their help with borehole temperature measurements in the field. We thank Tingting Ke, Ruyang Yu and Yi Li of Xi’an Jiaotong University for their assistance over various stages of this work.

6.3. A major geodynamic mechanism resulted in the thinning of the lithospheric mantle

References

Dong et al. (2014) indicated that hot material upwelling from the upper mantle started to rise northwards in the mid-to-late Mesozoic through the channel of the Inner Mongolia Suture Zone (Fig. 8a). Thus, the lithosphere experienced a certain degree thinning in the Erlian Basin, the Yinshan Block, and the Inner Mongolia Suture Zone. However, the lithosphere of the Ordos Block appears to be largely intact, based on seismic studies (Chen et al., 2009; Jiang et al., 2013; Tian et al., 2009). The thickness of the crust in study area is about the same as that of the Ordos Block. Therefore, we infer that the hot upwelling material only changed the lithospheric upper mantle. This is consistent with the result of the numerical simulation that some heat source invaded at the depth of 40−50 km. Santosh (2010) and Santosh et al. (2010) inferred that after the closure of the Paleo-Asian Ocean, the residual oceanic subduction slab in the NCC, and the asthenosphere convection below 300 km were the main driving force for the thinning of the lithosphere (Fig. 8b). Zhu et al. (2017) suggested that the driving force of the Mongolia - Baikal area rift was below 350 km. The upwelling mantle spread beneath Mongolia may be related to the deep subduction of the Pacific and the Indian Plates (Zhang et al., 2016). Mantle thermal convection is an important factor leading to the thinning of the lithosphere around the study area, whereas its tectonic evolution has experienced the closure of the Paleo-Asian Ocean, the subduction of the Pacific Plate (Kusky et al., 2014; Zhao, 2009), and the collision between the Indian and Eurasian Plates (Rowley, 1996). The geodynamic mechanism for lithospheric thinning is complex and a result of many factors. We have discussed only the most plausible mechanisms that could generate the heat flow anomaly in the Baiyinchagan Sag. However, the existing heat flow measurements are too sparse to define the extent of this anomaly and its thermotectonic significance, calling for further investigations.

Abdulagatova, Z., Abdulagatov, I., Emirov, V.N., 2009. Effect of temperature and pressure on the thermal conductivity of sandstone. Int. J. Rock Mech. Min. Sci. 46 (6), 1055–1071. Barkaoui, A.E., Correia, A., Zarhloule, Y., Rimi, A., Carneiro, J., Boughriba, M., Verdoya, M., 2013. Reconstruction of remote climate from borehole temperature measurement in the eastern part of Morocco. Clim. Change 118 (2), 431–441. Bayer, P., Rivera, J.A., Schweizer, D., Scharli, U., Blum, P., Rybach, L., 2016. Extracting past atmospheric warming and urban heating effects from borehole temperature profiles. Geothermics 64, 289–299. Bonnetti, C., Malartre, F., Huault, V., Cuney, M., Bourlange, S., Liu, X., Peng, Y., 2014. Sedimentology, stratigraphy and palynological occurrences of the late cretaceous erlian formation, Erlian Basin, Inner Mongolia, people’s Republic of China. Cretac. Res. 48, 177–192. Bredehoeft, J.D., Papaopulos, I.S., 1965. Rates of vertical groundwater movement estimated from the Earth’s thermal profile. Water Resour. Res. 1 (2), 325–328. Carslaw, H.S., Jaeger, J.C., 1959. Conduction of Heat in Solids, 2nd ed. Oxford University Press, New York, USA, pp. 510. Chen, L., Cheng, C., Wei, Z.G., 2009. Seismic evidence for significant lateral variations in lithospheric thickness beneath the central and western North China Craton. Earth Planet. Sci. Lett. 286 (1), 171–183. Clauser, C., Villinger, H., 1990. Analysis of conductive and convective heat transfer in a sedimentary basin, demonstrated for the Rheingraben. Geophys. J. Int. 100 (3), 393–414. Cui, Y.Q., Liu, X.H., Sun, Z.H., Yao, H., Han, C.Y., 2011. The deep physical geographic characteristics and Neopaleozoic geological structural exploration of Erlian basin, Inner Mongolia, china. Geological Bulletin of China 30 (2/3), 235–242 in Chinese with English abstract. Dong, H., Wei, W.B., Ye, G.F., Jin, S., Jones, A.G., Jing, J.E., Zhang, L.T., Xie, C.L., Zhang, F., Wang, H., 2014. Three‐dimensional electrical structure of the crust and upper mantle in Ordos Block and adjacent area: evidence of regional lithospheric modification. Geochem. Geophys. Geosystems 15 (6), 2414–2425. Gao, P., Qiu, Q.F., Jiang, G.Z., Zhang, C., Hu, S.B., Lei, Y.H., Wang, X.Z., 2018. Presentday geothermal characteristics of the Ordos Basin, western North China Craton: new findings from deep borehole steady-state temperature measurements. Geophys. J. Int. 214 (1), 254–264. Guo, P.Y., Niu, Y.L., Sun, P., Wang, X.H., Gong, H.M., Duan, M., Zhang, Y., Kong, J.J., Tian, L.Y., Wu, S.G., 2018. The Early Cretaceous bimodal volcanic suite from the Yinshan Block, western North China Craton: origin, process and geological significance. J. Asian Earth Sci. 160, 348–364. Guo, P.Y., Zhang, N., He, M.C., Bai, B.H., 2017. Effect of water saturation and temperature in the range of 193 to 373K on the thermal conductivity of sandstone. Tectonophysics 699, 121–128. Huang, J.L., Zhao, D.P., 2006. High‐resolution mantle tomography of China and surrounding regions. J. Geophys. Res. Solid Earth 111 (B9), 4813–4825. Huang, S.P., Pollack, H.N., Shen, P.Y., 2000. Temperature trends over the past five centuries reconstructed from borehole temperatures. Nature 403 (6771), 756–758. Huang, Z.X., Li, H.Q., Zheng, Y.J., Peng, Y.J., 2009. The lithosphere of North China Craton from surface wave tomography. Earth Planet. Sci. Lett. 288 (1), 164–173. Jiang, G.Z., Hu, S.B., Shi, Y.Z., Zhang, C., Wang, Z.T., Hu, D., 2019. Terrestrial heat flow of continental China: updated dataset and tectonic implications. Tectonophysics. Jiang, M.M., Ai, Y.S., Chen, L., Yang, Y.J., 2013. Local modification of the lithosphere beneath the central and western North China Craton:3-D constraints from Rayleigh wave tomography. Gondwana Res. 24, 849–864. Jiao, Y.X., Qiu, N.S., Li, W.Z., Zuo, Y.H., Que, Y.Q., Liu, F.L., 2013. The MesozoicCenozoic evolution of lithosphere thickness in the Ordos basin constrained by geothermal evidence. Chinese J. Geophys. 56 (9), 3051–3060 in Chinese with English abstract. Kusky, T.M., Windley, B.F., Wang, L., Wang, Z.S., Li, X.Y., Zhu, P.M., 2014. Flat slab subduction, trench suction, and craton destruction: comparison of the North China, Wyoming, and Brazilian cratons. Tectonophysics 630, 208–221. Mansure, A.J., Marshall, R.J., 1979. A vertical groundwater movement correction for heat flow. J. Geophys. Res. Atmos. 84 (B7), 3490–3496.

7. Conclusion Based on the above measurements and analysis, the following major conclusions can be drawn. (1) The range of the geothermal gradients in the Baiyinchagan Sag of the Erlian Basin in northern China is 36.6–45.4 ℃ km−1. (2) The rock thermal conductivities of the Saihantala, the Duhongmu, the Tenggeer and the Aershan Formations are found to be 1.15, 2.13, 1.99, and 2.15 W m−1 K-1, respectively. (3) Heat flow values range from 80.8 to 89.7 mW m−2 in the Baiyinchagan Sag, and the mean of the measured heat flow values is 84.9 ± 3.4 mW m−2, substantially higher than previously reported estimates based on oil-field formation testing data. The new measured high heat flow is consistent with the deep thermotectonic regime revealed by recent geophysical investigations. 10

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J. Zhang, et al. Pollack, H.N., Hurter, S.J., Johnson, J.R., 1993. Heat flow from the Earth’s interior: analysis of the global data set. Rev. Geophys. 31 (3), 267–280. Pollack, H.N., Huang, S.P., 2000. Climate reconstruction from subsurface temperatures. Annu. Rev. Earth Planet. Sci. 28, 339–365. Popov, Y.A., Pevzner, S.L., Pimenov, V.P., Romushkevich, R.A., 1999a. New geothermal data from the Kola superdeep well SG-3. Tectonophysics 306 (3), 345–366. Popov, Y.A., Pribnow, D.F., Sass, J.H., Williams, C.F., Burkhardt, H., 1999b. Characterization of rock thermal conductivity by high-resolution optical scanning. Geothermics 28 (2), 253–276. Powell, W.G., Chapman, D.S., Balling, N., Beck, A.E., 1988. Continental heat-flow density. In: Haenel, R., Rybach, L., Stegena, L. (Eds.), Handbook of Terrestrial Heat-Flow Density Determination. Spring, Netherlands, pp. 167–222. Ren, Z.L., Liu, C.Y., Feng, J.H., Pu, R.M., 2000. Research on geothermal history of Bayindulan depression in Erlian basin. Acta Geol. Sin. 21 (4), 42–45 in Chinese with English abstract. Rowley, D.B., 1996. Age of initiation of collision between India and Asia: a review of stratigraphic data. Earth Planet. Sci. Lett. 145 (1-4), 1–13. Safonova, I.Y., Seltmann, R., Kröner, A., Gladkochub, D., Kim, J.Y., Schulmann, K., Xiao, W.J., Komiya, T., Sun, M., 2011. A new concept of continental construction in the Central Asian Orogenic Belt: (compared to actualistic examples from the Western Pacific). Episodes 34 (3), 1–11. Santosh, M., 2010. Assembling North China Craton within the Columbia supercontinent: the role of double-sided subduction. Precambrian Res. 178 (1–4), 149–167. Santosh, M., Zhao, D.P., Kusky, T., 2010. Mantle dynamics of the Paleoproterozoic North China Craton: a perspective based on seismic tomography. J. Geodyn. 49 (1), 39–53. Sun, W.D., Ling, M.X., Wang, F.Y., Ding, X., Hu, Y.H., Zhou, J.B., Yang, X.Y., 2008. Pacific plate subduction and mesozoic geological event in Eastern China. Bulletin of Mineralogy, Petrology and Geochemistry 27 (3), 218–225 in Chinese with English abstract. Tang, J., Xu, W.L., Wang, F., Zhao, S., Wang, W., 2016. Early Mesozoic southward subduction history of the Mongol–okhotsk oceanic plate: evidence from geochronology and geochemistry of Early Mesozoic intrusive rocks in the Erguna Massif, NE China. Gondwana Res. 31, 218–240. Taniguchi, M., Shimada, J., Tanaka, T., Kayane, I., Sakura, Y., Shimano, Y., Dapaah‐Siakwan, S., Kawashima, S., 1999. Disturbances of temperature‐depth profiles due to surface climate change and subsurface water flow: 1. An effect of linear increase in surface temperature caused by global warming and urbanization in the Tokyo Metropolitan Area, Japan. Water Resour. Res. 35 (5), 1507–1518. Teng, J.W., Wang, F.Y., Zhao, W.Z., Zhang, Y.Q., 2010. Velocity structure of layered block and deep dynamic process in the lithosphere beneath the Yinshan orogenic belt and Ordos Basin. Chinese J. Geophys. 53 (1), 67–85 in Chinese with English abstract. Tian, X.B., Teng, J.W., Zhang, H.S., Zhang, Z.J., Zhang, Y.Q., Yang, H., Zhang, K.K., 2011. Structure of crust and upper mantle beneath the Ordos Block and the Yinshan Mountains revealed by receiver function analysis. Phys. Earth Planet. Inter. 184 (3–4), 186–193. Tian, Y., Zhao, D.P., Sun, R.M., Teng, J.W., 2009. Seismic imaging of the crust and upper mantle beneath the North China Craton. Phys. Earth Planet. Inter. 172 (3), 169–182. Uchida, Y., Sakura, Y., Taniguchi, M., 2003. Shallow subsurface thermal regimes in major plains in Japan with reference to recent surface warming. Phys. Chem. Earth 28 (9–11), 457–466. Wang, Q.S., Teng, J.W., An, Y.L., Zhang, Y.Q., 2010. Gravity field and deep crustal

structures of the Yinshan orogen and the northern Ordos basin. Progress in Geophys. 25 (5), 1590–1598 in Chinese with English abstract. Westaway, R., Younger, P.L., 2016. Unravelling the relative contributions of climate change and ground disturbance to subsurface temperature perturbations: case studies from Tyneside. UK. Geothermics 64, 490–515. Wilde, S.A., Zhou, J.B., 2015. The late Paleozoic to Mesozoic evolution of the eastern margin of the Central Asian Orogenic Belt in China. J. Asian Earth Sci. 113, 909–921. Xia, B., Thybo, H., Artemieva, I.M., 2017. Seismic crustal structure of the North China Craton and surrounding area: synthesis and analysis: NCC and surroundings’ crustal structure. J. Geophys. Res. Solid Earth 122, 5181–5207. Xiao, A.C., Yang, S.F., Chen, H.L., 2001. Geodynamic background on formation of Erlian basin. Oil Gas Geol. = Geol. Neft. I Gaza 22 (2), 137–140 in Chinese with English abstract. Xiao, W., Liu, Z., Du, J.H., Yi, S.W., 2004. Characteristic of geotherm-geopressure system in Erlian Basin. Xinjiang Petroleum Geology 25 (6), 610–613 in Chinese with English abstract. Xiong, S.Q., Yang, H., Ding, Y.Y., Li, Z.K., 2016. Characteristics of Chinese continent Curie point isotherm. Chinese J. Geophys. 59 (10), 3604–3617 in Chinese with English abstract. Xu, W., Huang, S.P., Zhang, J., Yu, R.Y., Zuo, Y.H., 2018. Present-day geothermal regime of the uliastai depression, Erlian Basin, North China. Energy Explor. Exploit. 0 (0), 1–17. Yang, Q.D., Wang, T., Guo, L., Tong, Y., Zhang, L., Zhang, J.J., Hou, Z.Q., 2017. Nd isotopic variation of Paleozoic–mesozoic granitoids from the Da Hinggan Mountains and adjacent areas, NE Asia: implications for the architecture and growth of continental crust. Lithos 272–273, 164–184. Zhai, M.G., Zhu, R.X., Liu, J.M., Meng, Q.R., Hou, Q.L., Hu, S.B., Liu, W., Li, Z., Zhang, H.F., Zhang, H.F., 2004. Time range of Mesozoic tectonic regime inversion in eastern North China Block. Sci. China Ser. D Earth Sci. 47 (2), 151–159. Zhang, F.X., Wu, Q.J., Grand, S.P., Li, Y.H., Gao, M.T., Demberel, S., Ulziibat, M., Sukhbaatar, U., 2016. Seismic velocity variations beneath central Mongolia: evidence for upper mantle plumes? Earth Planet. Sci. Lett. 459, 406–416. Zhang, S.H., Gao, R., Li, H.Y., Hou, H.S., Wu, H.C., Li, Q.S., Yang, K., Li, C., Li, W.H., Zhang, J.S., Yang, T.S., Keller, G.R., Liu, M., 2014. Crustal structures revealed from a deep seismic reflection profile across the Solonker suture zone of the Central Asian Orogenic Belt, northern China: an integrated interpretation. Tectonophysics 612–613, 26–39. Zhao, D.P., 2009. Multiscale seismic tomography and mantle dynamics. Gondwana Res. 15 (3–4), 297–323. Zhu, R.X., Chen, L., Wu, F.Y., Liu, J.L., 2011. Timing. Scale and mechanism of the destruction of the North China Craton. Sci China Earth Sci 54 (6), 789–797. Zhu, R.X., Xu, Y.G., Zhu, G., Zhang, H.F., Xia, Q.K., Zheng, T.Y., 2012. Destruction of the North China Craton. Sci China Earth Sci 55 (10), 1565–1587 in Chinese with English abstract. Zhu, W.P., Tian, W., Wei, C.J., Shao, J.A., Fu, B., Fanning, C.M., Chen, M.M., Wang, B., 2017. Late Paleozoic rift-related basalts from central Inner Mongolia, china. J. Asian Earth Sci. 144, 155–170. Zuo, Y.H., Wang, C.C., Tang, S.L., Hao, Q.Q., 2016. Mesozoic and Cenozoic thermal history and source rock thermal evolution of the Baiyinchagan sag, Erlian Basin, Northern China. J. Pet. Sci. Eng. 139, 171–184.

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