High elevation of Jiaolai Basin during the Late Cretaceous: Implication for the coastal mountains along the East Asian margin

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High elevation of Jiaolai Basin during the Late Cretaceous: Implication for the coastal mountains along the East Asian margin Laiming Zhang a,b,∗,1 , Chengshan Wang a,b,∗ , Ke Cao c , Qian Wang a,b,2 , Jie Tan a,b , Yuan Gao a,b a b c

State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Beijing 100083, China School of the Earth Science and Resources, China University of Geosciences, Beijing 100083, China The Key Laboratory of Marine Hydrocarbon Resources and Environment Geology, Qingdao Institute of Marine Geology, Qingdao 266071, China

a r t i c l e

i n f o

Article history: Received 20 April 2016 Received in revised form 14 September 2016 Accepted 21 September 2016 Available online xxxx Editor: H. Stoll Keywords: Late Cretaceous coastal mountains paleoelevation clumped isotopes East Asia Jiaolai Basin

a b s t r a c t A large body of evidence suggests that there were extensive coastal mountains along the East Asian margin during the Late Cretaceous. However, current knowledge of the paleo-mountains — the period, range, and elevation — is limited. Therefore, direct paleoaltimetry is needed to validate and evaluate the paleo-mountains in East Asia. Our study area is Jiaolai Basin, which is located at the East Asian continental margin. We estimate the paleoelevation of Jiaolai Basin during the Late Cretaceous using carbonate clumped isotope paleothermometry. After correcting for seasonal preference, latitudinal difference, and secular climate change, we conclude that the paleoelevation of Jiaolai Basin was almost certainly ≥2.0 km at ∼80 Ma. Combined with the evidence from stratigraphy, paleogeography, and paleoclimatology, our results suggest that the existence of coastal mountains along East Asia during the Late Cretaceous is likely and the model of Okhotomorsk–East Asia collision is preferred. © 2016 Elsevier B.V. All rights reserved.

1. Introduction During the Late Cretaceous, high-pressure metamorphic belts formed in the Taiwan, Japan, and the Sakhalin Islands (Fig. 1) (Isozaki et al., 2010). A great volume of volcanics and granites intruded along the East Asian continental margin (Yang, 2013). Simultaneously, the sedimentary basins in East Asia were uplifted and exhumed (Song et al., 2015). The regional hiatuses were widespread in these sedimentary basins, which extended from the continental margin to the hinterland (e.g. Nanhuabei Basin near Zhengzhou) (Li et al., 2014; Zhu et al., 2012). The evidence indicates that the orogeny during this period may have resulted in relatively extensive topographic changes along the East Asian margin. Based on structural geology and geochronological evidences for collision and orogenic exhumation from 100–89 Ma, Yang (2013) proposed coastal mountains extending from Southeast China to

*

Corresponding authors. E-mail addresses: [email protected] (L. Zhang), [email protected] (C. Wang). 1 Present address: School of Energy Resources, China University of Geosciences, Beijing 100083, China. 2 Present address: China Huadian Green Energy Co., Ltd., Beijing 100160, China. http://dx.doi.org/10.1016/j.epsl.2016.09.034 0012-821X/© 2016 Elsevier B.V. All rights reserved.

South Korea and Southwest Japan (Fig. 1). Two models have been proposed to explain the orogeny, such as the Pacific-induced compression (Song et al., 2015), or a collision between the Okhotomorsk Continental Block and the South China Block (Yang, 2013). Based on the thickness of denuded molasses accumulations, Chen (2000) suggested that the coastal mountains attained elevations of between 3500 m and 4000 m above sea level and width of 500 km from east to west. However, precise estimates of paleoelevations and the timing of maximum elevation have not yet been independently determined. Therefore, we propose a direct estimation of the paleoelevation from the East Asian continental margin to fully evaluate the topographic changes during the Late Cretaceous. Topography is a first-order expression of the buoyancy of the lithosphere and also strongly influences circulation of the atmosphere and global climate, therefore, paleoelevation change in the East Asian margin is one of the best available measures of Cretaceous continental dynamics and Cretaceous paleoclimate (Huntington and Lechler, 2015). However, given the conventional stable isotope paleoaltimetry, either based on temperatureelevation gradients or precipitation oxygen/hydrogen isotopeelevation gradients, it is difficult to unambiguously determine the paleoelevation (Huntington and Lechler, 2015; Peters et al., 2013). In practice, researchers have to assume one unknown (e.g., temperature) when estimating another unknown (e.g., δ 18 O value

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Fig. 1. Schematic geological map showing the key geological and geographical characteristics and the location of the study area (red star) during the Late Cretaceous. Yellow represents the piedmont basins of molasses sediments. Green represents the basins of dark-color sediments. Purple represents the basins of red-color sediments and brackish-saltwater sediments. The pink area shows the approximate range of the coastal mountains proposed by previous studies (Chen, 2000; Yang, 2013). JY = Jiayin Basin, SL = Songliao Basin, EL = Erlian Basin, JL = Jiaolai Basin, NH = Nanhuabei Basin, SB = Suibei Basin, HH = Nanhuanghai Basin, JH = Jianghan Basin, YM = Yuanma Basin, HY = Hengyang Basin. Modified after Song et al. (2015). (For interpretation of the references to color in this figure, the reader is referred to the web version of this article.)

of the water) because the carbonate formation temperature and δ 18 O value of the water from which it forms are interdependent (Kim and O’Neil, 1997). Moreover, the precipitation isotopic values (δ 18 O, δ 2 H, or δ D) are also controlled by the evaporation, moisture source, and seasonality of the mineral formation (Chamberlain and Poage, 2000). In this study, we use carbonate clumped isotope (47) thermometry to estimate the paleoelevation of Jiaolai Basin during the Late Cretaceous (Fig. 1). This method can reconstruct the growth temperatures of the carbonate minerals by evaluating the extent to which 13 C and 18 O are chemically bound to each other (clumped)

within the same carbonate ion group (Passey et al., 2010). It is based on a homogeneous isotope exchange equilibrium; therefore, the 47 temperatures are independent of the isotope compositions of the waters from which the carbonates grew (Passey et al., 2010). We first use carbonate clumped isotope (47) thermometry to obtain the paleotemperatures in Jiaolai Basin, Shandong Province from the Early Cretaceous to the Paleocene. We then calculate the temperature differences between Jiaolai Basin and the presumed low elevation site (Songliao Basin) during the Late Cretaceous using the same paleothermometer. Finally, we calculate the elevation based on the temperature-elevation lapse rate.

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Fig. 2. Stratigraphic columns showing the lithologies and samples for (A) the Early Cretaceous Laiyang Group (Jiaolai Basin), (B) the Late Cretaceous Wangshi Group (Jiaolai Basin), (C) the Paleocene Jiaozhou Formation (Jiaolai Basin), and (D) the Late Cretaceous Mingshui Formation (Songliao Basin). Pt = Proterozoic, Klx = Xiaoxianzhuang Formation, Klz = Zhifengzhuang Formation, Klm = Maershan Formation, Kq = Qingshan Group, Kwl = Linjiazhuang Formation, and Q = quaternary. Black arrows indicate the locations of paleosol carbonate samples for clumped isotope analyses and red arrows indicate the radiometric ages. (For interpretation of the references to color in this figure, the reader is referred to the web version of this article.)

2. Stratigraphy and sampling 2.1. Jiaolai Basin Jiaolai Basin is located in Shandong Peninsula in the eastern part of the North China Craton (Fig. 1). The basin is approximately 500 km long in the NE–SW direction and approximately 100 km wide in the E–W direction. It is a late Mesozoic extensional basin controlled by strike-slip and pull-apart activities of the NNEtrending faults (An et al., 2016). The basin consists of several E–W to WNW–ESE-trending grabens, which are filled with Cretaceous– Cenozoic fluvial–lacustrine sediments and volcanic rocks (Zhang et al., 2003). The paleolatitude of Jiaolai Basin during the Late Cretaceous is similar to where it is now (∼37◦ N) (Boucot et al., 2013). The oldest sedimentary cover within the basin is the Early Cretaceous Laiyang Group (Fig. 2), which is characterized by fluvial– lacustrine yellow to green sandstones, shales, conglomerates, and intercalated volcanic deposits. It comprises the Xiaoxianzhuang Formation (Klx), Zhifengzhuang Formation (Klz), Maershan Formation (Klm), Shuinan Formation (Kls), Longwangzhuang Formation (Kllw), and Qugezhuang Formation (Klq), from the bottom to the top. Above the Laiyang Group is the Early Cretaceous Qingshan Group, which is characterized by a series of intercalated volcanic and clastic rocks. It consists of four eruptive cycles: the Houkuang

Formation (Kqh), Bamudi Formation (Kqb), Shiquanzhuang Formation (Kqs), and Fanggezhuang Formation (Kqf). The sedimentary sequence is capped by an unconformity between the Lower Cretaceous and Upper Cretaceous, overlain by the Late Cretaceous Wangshi Group. The Wangshi Group is a red sedimentary rock series with purple sandy conglomerates, coarse sandstones, yellowish mudstones and siltstones of the lower Linjiazhuang (Kwl) and Xingezhuang (Kwx) formations, and brick-colored sandstoneconglomerates of the upper Hongtuya (Kwh) and Jingangkou (Kwj) formations (Zhang et al., 2003). In Jiaozhou city, the Late Cretaceous Hongtuya Formation (Kwh) is overlain by the Paleocene Jiaozhou Formation, which is overlain by Quaternary sediments. The Jiaozhou Formation is 1172 m thick and is characterized by grayish-green to grayish-purple mudstones in the lower part, and brick-red to red-purple and grayish-green to yellow-green siltstones/mudstone, and intercalated conglomerates and marlstones in the middle and upper parts (Y. Zhang et al., 2008). 2.1.1. Samples from the Lower Cretaceous We collected lacustrine carbonates from the Shuinan Formation of the Laiyang Group type section (120.85◦ N, 36.96◦ E). The section is approximately 7000 m long and the thickness of the strata is approximately 700 m (the burial depth of the section at least 3 km) (Fig. 2). The Shuinan Formation is composed of

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gray calcareous siltstones, gray-black shales, oil shales, and intercalated marlstones and limestones. This formation represents a humid deep-water lake environment (Zhang et al., 2003). Based on the biostratigraphy, the Shuinan Formation is Early Cretaceous in age. The zircon U–Pb and hornblende 40 Ar–39 Ar of the intercalated basaltic volcanic indicate that the samples are ∼130 Ma (Y. Zhang et al., 2008). 2.1.2. Samples from the Upper Cretaceous We collected paleosol carbonates from the Xingezhuang Formation of the Wangshi Group type section (120.73◦ N, 36.95◦ E) (Fig. 2). The Xingezhuang Formation is composed of purple, yellowgreen, and grayish-green fine sandstones to siltstones, and intercalated sandy conglomerates and marlstones (Zhang et al., 2003). The cyclothems composed by the fine sandstone–siltstone–mudstone are common in this formation. The paleosols are recognized at the top of the mudstones. This formation represents a semi-arid to arid shallow lake to fluvial environment. The paleosol carbonates (>30 cm below the paleosol surface) were collected from the top of the Xingezhuang Formation. Based on the biostratigraphy, the Hongtuya Formation and the top of the Xingezhuang Formation are Campanian in age (Fig. 2). The detrital zircon U–Pb age indicates that the bottom of the Hongtuya Formation and the top of the Xingezhuang Formation should be ≤77.3 Ma (An et al., 2016). The 40 Ar–39 Ar age of the intercalated volcanics at the top of the Hongtuya Formation is 73.5 ± 0.3 Ma (J. Zhang et al., 2008). The thickness of the Hongtuya Formation in this section is 448.56 m and the average deposition rate of the Wangshi Group is 73.2 m/million yrs (Y. Zhang et al., 2008). Therefore, the age of our samples from the Wangshi Group are approximately 79.63 ± 0.3 Ma and are most likely between 73.5 and 79.63 Ma. 2.1.3. Samples from the Paleocene We collected Paleocene carbonates from a borehole in Jiaozhou city (120.05◦ N, 36.27◦ E). The borehole is approximately 1600 m in depth and comprises the Paleocene Jiaozhou Formation and Upper Cretaceous strata. The paleosol carbonates (>30 cm below the paleosol surface) and lacustrine carbonates were collected from the cores taken between 250 m and 550 m (Fig. 2). Based on the biostratigraphy, the lower part of the Jiaozhou Formation is Late Cretaceous in age and the middle and upper parts of the Jiaozhou Formation are Paleocene in age (Y. Zhang et al., 2008). Therefore, these samples are Paleocene in age; however, the exact age of each sample is not well constrained. 2.2. The Songliao Basin The Songliao Basin is located in Northeast China and its paleolatitude was ∼45◦ N during the Late Cretaceous (Fig. 2) (Wang et al., 2013). The basin is filled predominantly with volcaniclastic, alluvial fan, fluvial and lacustrine sediments of the Late Jurassic, Cretaceous, and Paleogene ages (Wang et al., 2013). The Late Cretaceous stratigraphy was recovered in a borehole designated SK-In (north core) in the central part of Songliao Basin by the “Cretaceous Continental Scientific Drilling Program of China” (Wang et al., 2013). The paleosol carbonates (>30 cm below the paleosol surface) were collected in the Mingshui Formation. The Mingshui Formation is composed of gray-green, gray, black and brown-red shale and gray-green sandstone (Wang et al., 2013). The formation represents a semi-humid temperate environment during the waning phase of the basin’s life, when the fossil diversities decreased greatly in the number of classes and species compared to deeper formations, im-

plying an associated shrinkage of the basin and a suppression of the entire ecosystem (Wang et al., 2013). The five magnetozones (from the chron C32n to C29r, and the uppermost mixed polarity) identified in the Mingshui Formation indicate a Late Campanian to Early Danian age (Wang et al., 2013). The age of each sample can be calculated using the newly established astronomical time scale by tuning the identified 405-kyr eccentricity cycles to the La2010d astronomical solution (Wu et al., 2014) (see figures in Wu et al., 2014). 3. Method 3.1. Sample preparation For the paleosol carbonate nodules, the carbonate powder was drilled from the polished surfaces of the samples using a microdrill. In practice, the polished surfaces should be drilled no deeper than 2 mm to avoid unintentionally drilling the secondary carbonate portion. All of the carbonate powder drilled from different drill holes in each sample were mixed well. For the lacustrine carbonates, the samples were broken into small pieces using a hammer and gently powdered in an agate mortar. 3.2. Clumped isotope analyses Clumped isotope analyses (δ 13 C and δ 18 O were simultaneously obtained) were conducted at Johns Hopkins University (the lab now moved to University of Michigan, Ann Arbor) in three sessions following the method of Passey et al. (2010). The CO2 was liberated from 8 mg of pure carbonate powders (or equivalent impure carbonate powders) in an acid bath containing 100% H3 PO4 at 90 ◦ C for 10 min, then purified and introduced to a Thermo Scientific MAT 253 mass spectrometer using an automated system (Huntington et al., 2009; Passey et al., 2010). We report 47 values relative to the “absolute reference frame (ARF)” by periodically analyzing aliquots of enriched/depleted CO2 that were isotopically equilibrated at 30 ◦ C or 1000 ◦ C (Dennis et al., 2011). The δ 13 C and δ 18 O are reported relative to either the VPDB (mineral) or the VSMOW (water) scales. We also analyzed one of three internal carbonate standards, HAF Carrara, NBS-19, or 102-GC-AZ01, every day alongside samples to monitor system stability and precision. The 47 temperatures were calculated using the calibration of Passey and Henkes (2012) with an adjustment of 0.082h for the kinetic effects of calcite phosphoric (Passey and Henkes, 2012). The δ 18 O of soil water (δ 18 Owater ) was also calculated from the 47 temperatures and δ 18 O using the calibration of Kim and O’Neil (1997). In this study, the heated gas line (1000 ◦ C) and the equilibrium gas line (30 ◦ C) are consistent throughout the entire analyses (Table S1). The 48 vs. the heated gas was < 1h in magnitude for all of the samples (Table S1). The long-term averages are HAF Carrara (n = 10) 47 = 0.398 ± 0.020h and 102-GC-AZ01 (n = 10) 47 = 0.698 ± 0.014h (ARF, mean ±1σ standard deviation) for the first session; HAF Carrara (n = 5) 47 = 0.398 ± 0.005h, NBS-19 (n = 5) 47 = 0.405 ± 0.011h, and 102-GC-AZ01 (n = 9) 47 = 0.695 ± 0.007h (ARF, mean ±1σ standard deviation) for the second session; and HAF Carrara (n = 19) 47 = 0.392 ± 0.010h, NBS-19 (n = 21) 47 = 0.391 ± 0.011h, and 102-GCAZ01 (n = 35) 47 = 0.697 ± 0.012h (ARF, mean ±1σ standard deviation) for the third session (Table S1). 4. Results The clumped isotope thermometry provide an ideal method to reconstruct the paleoelevation, but the samples should be strictly collected and the results should be strictly assessed to assure the

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Table 1 Results of clumped isotope analyses. Sample ID

Age (Ma)

Na

δ 13 Cb (h, VPDB)

δ 18 Ob (h, VPDB)

Laiyang Group, Jiaolai Basin (Early Cretaceous, lacustrine carbonates) LY-06 ∼130 1 −2.008 −10.244 LY-19 ∼130 1 0.102 −7.072 SD-20 ∼130 1 −1.063 −2.325 SD-25 ∼130 2 −2.426 −8.148 SD-35 ∼130 2 −3.469 −7.315 SD-38 ∼130 1 −2.399 −17.417 SD-39 ∼130 1 −1.791 −18.801 Wangshi Group, Jiaolai Basin (Late Cretaceous, paleosol carbonates) JH-01b 79.63 3 −4.691 −6.600 JH-01c 79.63 3 −4.616 −6.613 JH-02c 79.63 2 −6.090 −6.899 JH-06b 79.63 2 −5.248 −8.507 JH-07 79.63 2 −5.934 −7.097 JH-12a 79.63 2 −6.620 −7.010 79.63 Averagef Jiaozhou Formation, Jiaolai Basin (Paleocene, paleosol carbonates and lacustrine carbonates) klc-002 Paleocene 2 −9.201 −7.895 klp03-01 Paleocene 2 −6.745 −7.017 klp06-01 Paleocene 2 −8.765 −7.627 klp07-01 Paleocene 2 −8.823 −7.339 klp11-01 Paleocene 2 −9.203 −8.189 klp13-01 Paleocene 2 −9.994 −5.866 klp14-01 Paleocene 2 −9.790 −5.502 klp15-01 Paleocene 2 −11.186 −7.108 klp17-01 Paleocene 2 −10.288 −7.023 klp20-01 Paleocene 2 −9.608 −7.004 f Paleocene Average Mingshui Formation, Songliao Basin (Late Cretaceous, paleosol carbonates) SK-54 70.08 2 −6.593 −13.832 SK-55 70.36 3 −7.051 −12.940 SK-28 71.05 2 −5.060 −11.949 SK-29 71.43 3 −7.868 −11.654 f 70.73 Average

47c (h, ARF)

T (47)d (◦ C)

δ 18 Owater e (h, VSMOW)

0.620 0.582 0.587 0.663 0.724 0.555 0.530

(0.013) (0.013) (0.013) (0.015) (0.015) (0.013) (0.013)

52.7 (5.9) 71.3 (7.0) 68.6 (6.9) 35.2 (6.0) 14.3 (4.7) 86.9 (8.0) 103.0 (9.1)

−2.8 (1.01) 3.4 (1.07) 7.8 (1.07) −3.8 (1.13) −7.2 (1.03) −4.8 (1.11) −4.1 (1.16)

0.703 0.702 0.713 0.675 0.713 0.706

(0.008) (0.012) (0.009) (0.009) (0.009) (0.009)

21.0 (2.6) 21.2 (4.1) 17.7 (3.0) 30.7 (3.5) 17.6 (3.0) 20.0 (3.1) 21.6 (4.9)

−5.1 (0.54) −5.1 (0.84) −6.1 (0.65) −5.1 (0.67) −6.3 (0.65) −5.7 (0.65)

0.674 0.656 0.674 0.645 0.653 0.687 0.660 0.658 0.661 0.674

(0.009) (0.018) (0.009) (0.009) (0.009) (0.011) (0.009) (0.009) (0.010) (0.014)

31.0 (3.5) 37.7 (7.1) 30.9 (3.5) 42.2 (3.9) 39.1 (3.8) 26.3 (4.0) 36.3 (3.7) 36.8 (3.7) 36.0 (4.0) 31.0 (5.1) 34.9 (4.7)

−4.4 (0.68) −2.2 (1.31) −4.1(0.68) −1.7 (0.70) −3.2 (0.69) −3.3 (0.80) −1.0 (0.69) −2.5 (0.69) −2.6 (0.76) −3.5 (1.00)

0.680 0.679 0.699 0.666

(0.009) (0.011) (0.009) (0.008)

28.6 29.0 22.3 34.0 28.4

−10.8 (0.67) −9.8 (0.80) −10.2 (0.65) −7.6 (0.56)

(3.4) (4.1) (3.2) (3.0) (4.6)

a

Number of unique analyses of CO2 from carbonate. Uncertainties on δ 13 C and δ 18 O are <0.07h and 0.06h respectively. c ARF = Absolute Reference Frame. Acid correction = 0.082h. Uncertainty is reported in parentheses. SE = SD/SQRT (N). If SD is less than the observed long-term SD of lab standards (0.013h), 0.013h is used. b

d e f

Calculated using the Equation (5) in Passey and Henkes (2012). Uncertainty is reported in parentheses. Calculated using the equation for calcite reported in Kim and O’Neil (1997). Uncertainty is reported in parentheses. The weighted-average temperature and error are calculated using the equations outlined in Huntington et al. (2009).

samples preserve unaltered climate signals, because of the carbonate ‘clumping’ is more vulnerable to diagenesis and metamorphism than many bulk compositional indices (Huntington and Lechler, 2015). 4.1. Jiaolai Basin For the Early Cretaceous Laiyang Group, the δ 13 C values range from −3.469h to 0.102h, and the δ 18 O values range from −18.801h to −2.325h. The clumped isotope compositions range from 0.530h to 0.724h with an average measurement error of 0.014h (1σ SE), corresponding to the 47 temperature range of 14.3 ◦ C to 103.0 ◦ C with an average measurement error of 6.8 ◦ C (1σ SE). The calculated δ 18 Owater ranges from −7.2h to 7.8h (Table 1; Fig. 3). Most 47 temperatures are above the Earthsurface temperature (Table 1). And the bulk isotope (δ 13 C and δ 18 O) compositions exhibit abnormal distributions. The optical and cathodoluminescence (CL) feathers (Fig. S1) show obvious speckled CL patterns of the microspars and spars in the fabric between the grains, therefore we conclude that the samples from the Laiyang Group were altered. Therefore, these samples will not be discussed later. For the Late Cretaceous Wangshi Group, the δ 13 C values range from −6.620h to −4.616h, and the δ 18 O values range from −8.507h to −6.600h. The clumped isotope compositions range from 0.675h to 0.713h with an average measurement error of

Fig. 3. The δ 18 Owater vs. 47 temperature in Jiaolai Basin. The gray solid circles represent the Early Cretaceous data, the black solid circles represent the Late Cretaceous data, and the open circles represent the Paleocene data. The gray arrows represent the Early Cretaceous data, which are beyond the scales. Contours of the δ 18 O (Cal) are calculated from Kim and O’Neil (1997).

0.009h (1σ SE), corresponding to the 47 temperature range of 17.6 ◦ C to 30.7 ◦ C with an average measurement error of 3.2 ◦ C (1σ SE). The calculated δ 18 Owater ranges from −6.3h to −5.1h (Table 1; Fig. 3). The optical and CL petrography of these nodules exhibits homogeneous dense micrite (Fig. S1) and the bulk isotopes (δ 18 O and δ 13 C) are reasonable. The correlation between the soil water δ 18 O (δ 18 Owater ) and 47 temperatures does not show sign of closed-system alteration (Fig. 3). All of the samples

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Table 2 The paleolatitudes, average ages, average temperatures and latitudinal differences for the Jiaolai Basin, Songliao Basin and Shuqualak.

Paleolatitude Average age (Ma) Average temp. (◦ C) Latitudinal difference 0.4 ◦ C/◦ Temp. difference

Jiaolai Basin

Songliao Basin

Shuqualak

∼37◦ N

∼45◦ N

∼35◦ N

79.63 21.6 ± 4.9a 0◦ 21.6 ± 4.9 /

70.73 28.4 ± 4.6a 8◦ 31.6 ± 4.7 10.0 ± 6.8c

79.52 25.9 ± 2.5b −2◦ 25.1 ± 2.5 3.5 ± 5.5d

a The average temperatures of the Jiaolai Basin and Songliao Basin were calculated using the carbonate clumped isotope paleothermometer. b The average temperature of the Shuqualak was calculated using the TEX86 paleothermometer (Linnert et al., 2014). c Between the Jiaolai Basin and Songliao Basin. d Between the Jiaolai Basin and Shuqualak.

erence of carbonate formation (see sections 5.1 and 5.5), or by secular climate change during the Late Cretaceous–Paleocene (see section 5.3). Therefore, because the paleotemperatures from Jiaolai Basin are not sufficient for the paleoelevation estimation, we still need more data from the presumed contemporary low elevation site. 4.2. Presumed low elevation site In this study, Songliao Basin is presumed to be a low elevation site during the Late Cretaceous. Songliao Basin experienced compression from the Pacific after ∼90 Ma, which is supported by the distinct unconformity T11 and zircon fission track data (Song et al., 2015). In this period, the stress from the east not only led to the westward migration of the depocenter but also resulted in the continuous uplift and erosion in the east (Wang et al., 2013). Afterwards, the extension regime prevailed in Songliao Basin and seawater incursion events were recognized in the basin (Hu et al., 2015). Therefore, Songliao Basin was thought to have been at a low elevation during the deposition of the Mingshui Formation. The four samples closest to 79.63 Ma were analyzed in this study. Their average age is 70.73 Ma, the calculated δ 18 Owater ranges from −10.8h to −7.6h, and their weighted-average temperature is 28.4 ± 4.6 ◦ C (Tables 1 and 2; Fig. 4). 5. Discussion

Fig. 4. The paleotemperatures of Songliao Basin, Jiaolai Basin and Shuqualak. The green diamonds are 47 temperatures from Songliao Basin, the purple triangles are 47 temperatures from Jiaolai Basin, and the yellow squares are TEX86 temperatures from Shuqualak. The larger symbols represent the average ages and temperatures. The errors of the paleotemperatures are represented by a horizontal bar. (For interpretation of the references to color in this figure, the reader is referred to the web version of this article.)

are ≤100 Ma and the burial depths are no more than 2.0 km (Y. Zhang et al., 2008), thus the samples at most only slightly influenced by solid-state C–O bond reordering (Henkes et al., 2014). All of the samples were collected within 10 m in thickness, their average age is 79.63 Ma, and the weighted-average temperature is 21.6 ± 4.9 ◦ C (Tables 1 and 2; Fig. 4). The weighted-average temperature and error are calculated using the equations outlined in Huntington et al. (2009). For the Paleocene Jiaozhou Formation, the δ 13 C values range from −11.186h to −6.745h, and the δ 18 O values range from −8.189h to −5.502h. The clumped isotope compositions range from 0.645h to 0.687h with an average measurement error of 0.011h (1σ SE), corresponding to the 47 temperature range of 26.3 ◦ C to 42.2 ◦ C with an average measurement error of 4.2 ◦ C (1σ SE). The calculated δ 18 Owater ranges from −4.4h to −1.0h (Table 1; Fig. 3). The optical and CL petrography of these nodules exhibits homogeneous dense micrite (Fig. S1) and the bulk isotopes (δ 18 O and δ 13 C) are reasonable. The correlation between the soil water δ 18 O (δ 18 Owater ) and 47 temperatures does not show sign of closed-system alteration (Fig. 3). All of the samples are ≤70 Ma and the burial depths are no more than 1.0 km (Y. Zhang et al., 2008), thus the samples at most only slightly influenced by solidstate C–O bond reordering (Henkes et al., 2014). These samples are Paleocene in age; however, the exact age of each sample is not well constrained. Their weighted-average temperature is 34.9 ± 4.7 ◦ C (Table 1). In Jiaolai Basin, we determined that the average temperature during the Late Cretaceous is much lower (∼15 ◦ C) than that during the Paleocene (Fig. 3). This temperature difference may either have been caused by elevation differences, by the seasonal pref-

We estimate the paleoelevation of Jiaolai Basin during the Late Cretaceous according the approach outlined by Snell et al. (2014). Several factors should be considered in the paleoelevation estimation, which may introduce errors to the result: the seasonal preference of the paleotemperature proxies, the differences in the paleolatitudes between each site, the secular climate change during the temporal gap between the samples from Jiaolai Basin and the presumed low elevation site, and the different choices in lapse rates. 5.1. Seasonal preference of paleotemperature proxies The 47 temperature has a seasonal preference that depends on the season in which the carbonate nodule formed (Burgener et al., 2016; Quade et al., 2013; Ringham et al., 2016). In previous studies, the majority of soil carbonates were summer biased (Hough et al., 2014; Passey et al., 2010; Quade et al., 2013; Snell et al., 2014); although, there are still a few that formed in other seasons (Burgener et al., 2016; Peters et al., 2013; Ringham et al., 2016). Previous studies also confirmed that the monsoon or monsoonal-like climate prevailed in East Asia since the Latest Cretaceous (Chen et al., 2013; Quan et al., 2014), and there were no major changes in the climate pattern during the Late Cretaceous to Paleogene. In a monsoon climate, the soil carbonate likely formed in the dry half-cycles of the warmest months (Quade et al., 2013), immediately before the cooling effects of the monsoon rains and after the hottest part of the summer (Chen et al., 2013). The lacustrine carbonates also appear to be biased toward formation during the warm season (Huntington and Lechler, 2015; Huntington et al., 2010). With respect to Songliao Basin, the 47 temperatures (22– 34 ◦ C) are generally the same as the coeval 47 temperatures of fossil shells (23–31 ◦ C) and ∼15 ◦ C higher than the leaf range through the mean annual temperature from the similar paleolatitudes in North America (Tobin et al., 2014). Therefore, the carbonate nodules in Songliao Basin were concluded to have formed during the summer. With respect to Jiaolai Basin, the Paleocene samples are clearly summer biased because almost all of the 47 temperatures from

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the paleosol carbonates are higher than the warm-season biased 47 temperature from the lacustrine carbonate (klc-002). In addition, the samples temperatures are much higher than the Paleocene marine paleotemperature record (Friedrich et al., 2012; Linnert et al., 2014). Because there were no major changes in the climate pattern during the Late Cretaceous to Paleogene (Chen et al., 2013; Quan et al., 2014), and other Late Cretaceous to Paleogene samples from both Songliao Basin and Jiaolai Basin were summer biased, we infer that the Late Cretaceous samples from Jiaolai Basin formed during the summer (see section 5.4). 5.2. Correcting for latitudinal difference To account for the effect of latitudinal difference between Jiaolai Basin and the presumed low elevation site (8◦ for Songliao Basin) on the paleoelevation estimation, we applied a correction to the average paleotemperature of Songliao Basin using the latitudinal temperature gradient during the Late Cretaceous (Pucéat et al., 2007). The paleotemperature of Songliao Basin is revised to be at the same paleolatitude of Jiaolai Basin (∼37◦ N). Both the terrestrial and marine latitudinal temperature gradients during the Late Cretaceous are 0.4 ◦ C/◦ latitude from 30◦ N to 50◦ N (Pucéat et al., 2007). Accordingly, we revise the Songliao average temperature from 28.4 ± 4.6 ◦ C to 31.6 ± 4.7 ◦ C (Table 2). 5.3. Correcting for secular climate change After the latitudinal correction, the temperature difference between Jiaolai Basin and Songliao Basin is 10.0 ◦ C, which still needs to be corrected for the secular climate changes during the ∼9 Ma gap between the average age of samples from Jiaolai Basin and Songliao Basin. Unfortunately, a high-resolution Late Cretaceous terrestrial paleotemperature record in East Asia is not available. We have to use the δ 18 O record of the benthic foraminifera from the Pacific Ocean (Friedrich et al., 2012) through the CampanianMaastrichtian. Their ages are adjusted to “The Geologic Time Scale 2012” (Gradstein et al., 2012). To justify the correction for climate change by benthic δ 18 O record, the global climate record should generally reflect the paleoclimate evolution of Songliao Basin during the Late Cretaceous, both the directions and the amplifications of the climate changes. The Late Cretaceous paleoclimate record in Songliao Basin was established based on the pollen/spore data (Wang et al., 2013) and δ 18 O of the ostracods (Wang et al., 2013). The paleoclimate records are comparable between Songliao Basin and the Cretaceous oceans (Friedrich et al., 2012; Wang et al., 2013), although all of the climate changes in Songliao Basin generally lag behind the changes in the Cretaceous oceans for 1–2 Ma. The mismatches may be introduced by the uncertainties in dating or the slower responses to the global climate changes by the terrestrial system (Wang et al., 2013). Nevertheless, these mismatches can be considered in the corrections and will not affect the final results. Additionally, the amplifications of climate changes at Songliao Basin relative to the global changes are small (<2 ◦ C). Therefore, we conclude the global climate record (benthic δ 18 O) broadly reflect the paleoclimate evolution in Songliao Basin during the Late Cretaceous (Wang et al., 2013). Thus, the benthic δ 18 O data can be applied to the corrections. In the ∼9 Ma gap (79.63 for Jiaolai Basin and 70.73 for Songliao Basin), the change in the benthic δ 18 O is 0.4h in the Pacific Ocean (Fig. 5). Based on the relationship of temperature–δ 18 O between water and calcite of ∼4.8 ◦ C/h (Bemis et al., 1998), the corresponding temperature difference is 1.9 ◦ C (Fig. 5). Factoring for the 1.9 ◦ C cooling, the temperature difference should be 11.9 ◦ C. Due to the mismatches in climate change between the terrestrial and marine records, and the uncertainties of the sample ages

Fig. 5. Plot of temperature difference versus elevation. Left column: The gray symbols represent the benthic foraminiferal δ 18 O data from the Pacific Ocean and the black curve represents the average trends during the Campanian–Maastrichtian (Friedrich et al., 2012). The green arrow indicates the δ 18 O increase (cooling) in the marine record according to the average ages. The blue arrow indicates the maximum δ 18 O increase (cooling) in the marine record. The red arrow indicates the maximum δ 18 O decrease (warming) in the marine record. Right column: 10.0 ◦ C is the temperature difference between Jiaolai Basin and Songliao Basin after the latitudinal correction. The colored vertical lines represent the ranges of elevation estimates according to different benthic foraminiferal δ 18 O changes. The gray shaded region indicates the range lapse rate ranges of 4.0–6.0 ◦ C/km. Modified from Snell et al. (2014). (For interpretation of the references to color in this figure, the reader is referred to the web version of this article.)

(especially the Jiaolai samples), the possible maximum warming and cooling during the temporal gap are also estimated. The maximum warming is 1.8 ◦ C and the maximum cooling is 4.4 ◦ C (Fig. 5), corresponding to the temperature differences of 9.2 ◦ C and 14.4 ◦ C. 5.4. Lapse rates and paleoelevation estimation According to Wolfe (1992), the global atmospheric lapse rate (6 ◦ C/km) is generally higher than the lapse rates from the soil temperature and lake surface temperature. The lapse rates calculated by the 47 temperature of modern soil carbonates from the Himalayas (Quade et al., 2013) and central Rocky Mountains (Hough et al., 2014) are 4.0 ◦ C/km and 3.6 ◦ C/km, respectively. And the lapse rate calculated by the 47 temperature of modern lake carbonates from the Colorado Plateau (Huntington et al., 2010) is 4.2 ◦ C/km. The lapse rates also change with seasonality and region. The regional atmospheric lapse rates in China for different seasons are 3.75–5.15 ◦ C/km (Meyer, 1992), and the regional atmospheric lapse rates in East China for different distances from the eastern and southeastern coast are 4.0–5.0 ◦ C/km (Fang and Yoda, 1988). Therefore, we suggest a lapse rate of 4.0–6.0 ◦ C/km as being most plausible for the study area (Fig. 5). By using this lapse rate range, the paleoelevation of Jiaolai Basin was 2.0–3.0 km during the Late Cretaceous (Fig. 5). Considering the uncertainty of the temperature differences (9.2–14.4 ◦ C), the paleoelevation range extends to 1.5–3.6 km (Fig. 5). Because the Late Cretaceous is under a “greenhouse” state, the lapse rate would be lower than at present. Thus, the estimated paleoelevations are underestimated and should be considered to be minimum values (Snell et al., 2014). We believe the most reasonable paleoelevation in Jiaolai Basin during the Late Cretaceous was ≥2 km.

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Table 3 Evaluate the influence of seasonal preference on paleoaltimetry. Situation

Summer (+10–15 ◦ C)b (1) +10 ◦ C (2) +15 ◦ C

Temperature difference (◦ C)

Elevation (km) 4–6 ◦ C/km

3.5 ± 5.5a 13.5 18.5

Other seasons (+3.5–5 ◦ C)c (1) +3.5 ◦ C 7.0 (2) +5 ◦ C 8.5

2.3–3.4 3.1–4.6

1.2–1.8 1.4–2.1

a The average temperature difference between the Jiaolai Basin and Shuqualak after correction for latitudinal difference. b The carbonate nodules formed in summer and the 47 temperature generally exceeds the mean annual air temperature by 10–15 ◦ C on average due to the seasonal preference and solar heating. c The carbonate nodules formed not in summer and the 47 temperature generally exceeds the mean annual air temperature by 3.5–5.0 ◦ C on average due to solar heating.

5.5. Evaluation of the influence of seasonal preference We concluded that the carbonate nodules in Jiaolai Basin during the Late Cretaceous were formed in the summer and their 47 temperatures generally exceed the mean annual air temperature by 10–15 ◦ C on average. However, we cannot completely rule out the possibility that these samples formed in other seasons. If these paleosol carbonates were formed in the spring or autumn, the corrected temperature difference between Songliao Basin and Jiaolai Basin resulted not only in elevation differences but also seasonal temperature differences (Burgener et al., 2016; Peters et al., 2013; Ringham et al., 2016). Therefore, we elected to evaluate the influence of seasonal preference on paleoaltimetry. Two situations are considered: 1) The Late Cretaceous carbonate nodules in Jiaolai Basin formed in the summer and their 47 temperature generally exceeds the mean annual air temperature by 10–15 ◦ C on average due to the seasonal preference and solar heating (Hough et al., 2014; Quade et al., 2013); 2) the Late Cretaceous carbonate nodules in Jiaolai Basin formed in other seasons and their 47 temperature generally exceeds the mean annual air temperature by 3.5–5.0 ◦ C due to solar heating (Hough et al., 2014; Peters et al., 2013; Quade et al., 2013). Here, Shuqualak is selected as the presumed low elevation site. Shuqualak is located in Mississippi, USA and its paleolatitude is ∼35◦ N during the Late Cretaceous (Linnert et al., 2014). The mean annual sea surface temperatures were calculated using a TEX86 palaeothermometer (Linnert et al., 2014). The calibration used in this study is TEXL86 (Kim et al., 2010). The five data points closest to 79.63 Ma are used in this study. The average age is 79.52 Ma, and the average temperature is 25.9 ± 2.5 ◦ C (Table 2; Fig. 4). After considering the latitudinal difference, the temperature difference between Jiaolai Basin and Shuqualak is 3.5 ◦ C (Table 2). By using a lapse rate of 4.0–6.0 ◦ C/km, the paleoelevation of Jiaolai Basin was 2.3–4.6 km for situation 1 and 1.2–2.1 km for situation 2 (Table 3). The summer-biased situation agrees well with the result estimated in the section above. We also observed that even if the Late Cretaceous paleosol carbonates in Jiaolai Basin formed in other seasons, the paleoelevation of Jiaolai Basin would still be relatively high during the Late Cretaceous. 6. Implications 6.1. The uplift of the Jiaolai Basin The Jiaolai Basin experienced intense WNW-ESE extension during the Early Cretaceous, which led to development of widespread

volcanism, normal faulting and basin subsidence (Zhang et al., 2003). However, during the late Early Cretaceous (∼100 Ma), the tectonic stress regime changed dramatically to NW–SE compression (Zhang et al., 2003). In addition, Lee et al. (2011) found that the southern end of the Tan-Lu fault system (including Jiaolai basin) may have been under a trans-compressive stress-field environment at 100–80 Ma. The unconformity between the Late Cretaceous Wangshi Group and the underlying late Early Cretaceous Qingshan Group indicates that the Jiaolai Basin was uplifted in this period (Zhang et al., 2003). Therefore, we speculated the Jiaolai Basin was uplifted and exhumed at ∼100 Ma caused by this sudden NW–SE shortening event. This is also consistent with evidences from the apatite fission track ages of the Linglong granites in the northeastern part of Jiaolai Basin and the zircon fission track ages of the Luxi granites in the western part of Jiaolai Basin (Song et al., 2015). 6.2. Evidences for the coastal mountains along the East Asian continental margin Our results broadly support the existence of coastal mountains during the Late Cretaceous. The Jiaolai Basin located in the margin of the coastal mountains and its paleoelevation was certainly ≥2 km at ∼80 Ma (the uncertainty of the age see section 2.1.2), which means the paleoelevation of the coastal mountains should be much higher than 2 km during the climax of compression (∼90 Ma). In addition to the evidence derived from the paleoaltimetry, many stratigraphic, paleogeographic, and paleoclimatic studies have also suggested that large-scale mountains were rapidly uplifted along the East Asian continental margin during the same period. Firstly, during the late Early Cretaceous (∼100 Ma), the sudden NW–SE shortening event occurred in the whole East China (Li et al., 2014; Zhang et al., 2003; Song et al., 2015). Meanwhile, a major unconformity occurred in whole East China during the end of the Early Cretaceous (Zhu et al., 2012; Li et al., 2014), which indicates major mountains and basins were rapidly uplifted and exhumed between the Lower and Upper Cretaceous boundary. Secondly, conglomerates were widely distributed along the coast of East China since the Cenomanian, including the whole Southeast China, Subei Basin, Jiaolai Basin, and Nanhuabei Basin, representing the molasses sediments in piedmont basins (Fig. 1) (Chen, 2000). Simultaneously, coastal retreat from Southeastern China was also recognized (Hu et al., 2012). Thirdly, a paleoclimatic differentiation between Northeast China and the rest of East China indicates that the moisture from the oceans into the hinterland during the Late Cretaceous seems to have been blocked (Boucot et al., 2013; Chen, 2000; Hasegawa et al., 2012) (Fig. 1). In Northeast China, widespread coal and oil shale indicate a humid climate. In the remainder of East China, widespread evaporites, red beds, and deserts indicate the presence of a semi-arid to arid climate (Boucot et al., 2013; Hasegawa et al., 2012). All these evidences indicate that extensive topographic changes (coastal mountains) along the East Asian continental margin were likely. 6.3. The tectonic model There are two models proposed for the uplift of the coastal mountains. One model suggests that the East Asian margin experienced successive oceanic subduction and the coastal mountains were uplifted by the subduction of Pacific-Izanagi ridge (Charvet, 2013; Isozaki et al., 2010). The geophysical data indicate that the paleo-Pacific slab was subducted as far as below the Taihang Mountains, suggesting the long-term oceanic subduction beneath Eurasian continent (Isozaki et al., 2010; Zhu et al., 2015), which leads to widespread Early Cretaceous bimodal volcanic rocks in

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Fig. 6. Simplified cartoon showing Cretaceous tectonic evolution in East China. (A) Early Cretaceous (135–100 Ma); (B) late Early Cretaceous–early Late Cretaceous (100–85 Ma); (C) Late Cretaceous (85–66 Ma). The principal extension/compression direction in the East China and plate motion direction in the Pacific Ocean from Zhu et al. (2012).

the Northeast China (Xu et al., 2013). In addition, during the Cretaceous, the principal extension/compression direction of the East China was perpendicular to the subduction trench of the paleoPacific Plate, and parallel to the subduction direction of the paleoPacific Plate (Fig. 6) (Zhang et al., 2003; Zhu et al., 2012, 2015). The acceleration of the drift rates of the Izanagi plate corresponds with the peak time (87–89 Ma) of HP metamorphism in the Cretaceous accretion units of Taiwan, Japan and Sakhalin islands (Song et al., 2015). However, several lines of evidence contradict the successive oceanic subduction model and suggest the thrusting of a microcontinent is likely needed (Charvet, 2013) which can only be explained by the continental collision model. Yang (2013) suggests the uplift may be caused by a collision between the Okhotomorsk Continental Block and the South China Block at approximately 100 Ma. The Okhotomorsk Continental Block, currently residing be-

low the Okhotsk Sea in Northeast Asia, was located in the interior of the Izanagi Plate before the Late Cretaceous. The arguments are: Firstly, current oceanic subduction model is unable to build the ubiquitous strike-slip features with the sub-horizontal nappes in SW Japan (Charvet, 2013; Yang, 2013), whereas will induce an acceleration of the tectonic erosion and collapse of the upper plate. The geometry is actually similar to the one encountered in collisional orogeny. Secondly, the oceanic subduction model is unable to explain the peak metamorphic pressure of 2.9 to 3.8 GPa (accretionary-type subduction scenario vary from 0.7 to 2.0 GPa) of the Sanbagawa Belt in SW Japan (Charvet, 2013; Yang, 2013). In addition, the fast exhumation of at least 2.5 cm/yr of the Sanbagawa Belt during 89–85 Ma could not be achieved in the accretionary-type subduction setting (range between 1 and 5 mm/yr) (Charvet, 2013; Yang, 2013).

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Thirdly, the oceanic subduction model can’t explain the magmatic hiatus during the early Late Cretaceous (100–90 Ma) (Xu et al., 2013; Zhang et al., 2014), when few igneous rocks can be found in East China. The Okhotomorsk–East Asia collision proposed here provides an interpretation for this event. Fourthly, the oceanic subduction model can’t explain the sudden switch of tectonic stress regime in East Asia including the Jiaolai Basin at the end of the Early Cretaceous. This event not only affected the East Asian continental margin, it also had significant effects in the East Gobi Basin in Mongolia, a place far to the plate margin. In the East Gobi Basin, a brief period of compressional and/or transpressional basin inversion occurred at the end of the Early Cretaceous (Graham et al., 2001). With respect to the evidences in this study, although the subduction of ocean floor relief can uplift the land surface to form local coastal mountains (von Huene and Ranero, 2009), It is contradict with the extensive coastal mountains along East Asian margin extended from Southeast China to Southwest Japan and South Korea (Yang, 2013). Along the Americas, only the subductions of the Yakutat terrane in central Alaska and Cocos Ridge in central America were clearly associated with exceptionally high coastal mountains (>5000 m and >3000 m respectively), others only produced modest local coastal uplifts and minor permanent deformations. However, our results suggest the paleoelevation of the coastal mountains should be much higher than 2 km, between 3500 m and 4000 m above sea level, during the climax of compression. Therefore, our new data prefer the model of Okhotomorsk–East Asia collision to be the cause of the uplift of the coastal mountains. We should note that this study only provides a paleoelevation estimation for a single window of time (Late Cretaceous). Further constraint of the paleo-mountains will require more paleoelevation estimations along the coast of East Asia during the Cretaceous.

During this period, a continental-scale NW–SE shortening event occurred in East China. Major mountains and basins (e.g. the Jiaolai Basin) were rapidly uplifted and exhumed along East Asia margin (Zhang et al., 2003) and in the hinterland of East Asia (e.g. East Gobi Basin in Mongolia) (Graham et al., 2001). The uplift of northeast-striking mountain ranges in East Asia results in the deposition of red clastic sediments and eolian sands in China and Mongolia during the Late Cretaceous (Boucot et al., 2013; Hasegawa et al., 2012). 6.4.3. The Late Cretaceous (∼85–66 Ma) crustal extension When the southwestern end of the Okhotomorsk Block passed the NE-striking Asian margin, the space left by the departed lithosphere of the Okhotomorsk was rapidly filled by subduction of oceanic lithosphere to its south (Yang, 2013). The relaying Pacific Plate moved generally northwards from the Late Cretaceous to Paleogene (Zhu et al., 2012). The principal extension direction changed to nearly N–S in the Late Cretaceous-Paleogene (Zhu et al., 2012), which caused a new phase of regional crustal subsidence along the youngest E–W trending extensional structures (Zhang et al., 2003), and large-scale Late Cretaceous-Paleogene red-colored sedimentary basins (e.g. Wangshi Group in the Jiaolai Basin). The calculated paleoelevation of the Jiaolai Basin was almost certainly ≥2 km at ∼80 Ma (Fig. 6C). This stage is characterized by the alkaline and cal-alkaline basalts (Zhang et al., 2014). Their geochemical and isotopic characteristics indicate that they were completely different from those of the Early Cretaceous one, but similar to those of back-arc basalts from the Japan Sea, which thus provides a petrological evidence for the contribution of subducted Pacific slab to the Late Cretaceous magmatism in East China (J. Zhang et al., 2008). 7. Conclusion

6.4. The tectonic evolution of East China during the Cretaceous 6.4.1. The Early Cretaceous (∼135–100 Ma) crustal extension During the Early Cretaceous, the Izanagi Plate and its inclusion Okhotomorsk Continental Block together drifted nearly orthogonally (WNW-wards) to the East Asian margin (Zhu et al., 2012). After the final closure of the Mongol-Okhotsk Ocean, East China switched rapidly to a NW–SE extension regime, which was attributed to a far-field, back-arc extension related to roll-back of the trench in the Izanagi Plate (Fig. 6A) (Xu et al., 2013; Zhu et al., 2015). During this period, widespread magmatic rocks and extensional sedimentary basins of South China exhibit a pronounced NE-striking distribution (Li et al., 2014). Meanwhile, the intensive development of Early Cretaceous magmatism, extensional deformation and associated gold mineralization in North China, indicate that the eastern North China Craton was destroyed during this period (Zhang et al., 2014). The Jiaolai Basin was formed and characterized by widespread silicic to intermediate volcanism, normal faulting and basin subsidence (Zhang et al., 2003). 6.4.2. The late Early Cretaceous–early Late Cretaceous (∼100–85 Ma) collision At about 100–85 Ma (Fig. 6B), the WNW-wards-moving Okhotomorsk Block collided with the East Asian margin (∼100 Ma) then began to move northward with the NNW-wards-moving Izanagi Plate (∼90 Ma) (Yang, 2013). This collision event ended the large-scale volcanic-intrusive in East China and formed the maximum of high-pressure metamorphism in the Cretaceous accretion units of Taiwan, Japan and Sakhalin islands (Song et al., 2015; Yang, 2013). It also caused the formation of a sinistral strike-slip fault system in the South China Block and a dextral strike-slip fault system from Japan to North China (Yang, 2013).

In this study, we estimated the paleoelevation of Jiaolai Basin during the Late Cretaceous based on the clumped isotopes of paleosol carbonates. The conclusions are: (1) After correcting for seasonal preference, latitudinal difference, and secular climate changes, we determined that the paleoelevation of Jiaolai Basin was 2.0–3.0 km during the Late Cretaceous (with uncertainties of 1.5–3.6 km). We believe that the paleoelevation was almost certainly ≥2 km at ∼80 Ma. (2) Our results support the hypothesis of coastal mountains along East Asia during the Late Cretaceous. We suggest that the coastal mountains along the East Asian continental margin were uplifted at least since ∼100 Ma, and its paleoelevation should be much higher than 2 km during the climax of compression (∼90 Ma). (3) Our data prefer the model of Okhotomorsk–East Asia collision to be the cause of the uplift of the coastal mountains. Acknowledgements We thank E. Yu for assistance in the field. We would thank S. Liu, who reviewed our manuscript and gave us many useful comments. We would thank Benjamin H. Passey for the use of the laboratory facilities and for performing a quality assurance of the data. Laiming Zhang is supported by a scholarship by the Chinese Scholarship Council (201406400033). This study was financially supported by the National Basic Research Program of China (973 Project: 2012CB822000), the National Natural Science Foundation of China (41402105), and the Fundamental Research Funds for the Central Universities (53200859417).

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Appendix A. Supplementary material Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2016.09.034. References An, W., Kuang, H.W., Liu, Y.Q., Peng, N., Xu, K.M., Xu, H., Zhang, P., Wang, K.B., Chen, S.Q., Zhang, Y.X., 2016. Detrital zircon dating and tracing the provenance of dinosaur bone beds from the Late Cretaceous Wangshi Group in Zhucheng, Shandong, East China. J. Palaeogeogr. 5, 72–99. Bemis, B.E., Spero, H.J., Bijma, J., Lea, D.W., 1998. Reevaluation of the oxygen isotopic composition of planktonic foraminifera: experimental results and revised paleotemperature equations. Paleoceanography 13, 150–160. Boucot, A.J., Shen, X., Scotese, C.R., Morley, R.J., 2013. Phanerozoic Paleoclimate: An Atlas of Lithologic Indicators of Climate. Society for Sedimentary Geology, Tulsa, Oklahoma. Burgener, L., Huntington, K.W., Hoke, G.D., Schauer, A., Ringham, M.C., Latorre, C., Díaz, F.P., 2016. Variations in soil carbonate formation and seasonal bias over >4 km of relief in the western Andes (30◦ S) revealed by clumped isotope thermometry. Earth Planet. Sci. Lett. 441, 188–199. Chamberlain, C.P., Poage, M.A., 2000. Reconstructing the paleotopography of mountain belts from the isotopic composition of authigenic minerals. Geology 28, 115–118. Charvet, J., 2013. Late Paleozoic–Mesozoic tectonic evolution of SW Japan: a review – reappraisal of the accretionary orogeny and revalidation of the collisional model. J. Asian Earth Sci. 72, 88–101. Chen, J., Zhao, P., Wang, C., Huang, Y., Cao, K., 2013. Modeling East Asian climate and impacts of atmospheric CO2 concentration during the Late Cretaceous (66 Ma). Palaeogeogr. Palaeoclimatol. Palaeoecol. 385, 190–201. Chen, P.J., 2000. Paleoenvironmental changes during the Cretaceous in eastern China. In: Hakuyu, O., Nlall, J.M. (Eds.), Developments in Palaeontology and Stratigraphy. Elsevier, pp. 81–90. Dennis, K.J., Affek, H.P., Passey, B.H., Schrag, D.P., Eiler, J.M., 2011. Defining an absolute reference frame for ‘clumped’ isotope studies of CO2 . Geochim. Cosmochim. Acta 75, 7117–7131. Fang, J.Y., Yoda, K., 1988. Climate and vegetation in China (I). Changes in the altitudinal lapse rate of temperature and distribution of sea level temperature. Ecol. Res. 3, 37–51 (in Chinese with English abstract). Friedrich, O., Norris, R.D., Erbacher, J., 2012. Evolution of middle to Late Cretaceous oceans—a 55 m.y. record of Earth’s temperature and carbon cycle. Geology 40, 107–110. Gradstein, F.M., Ogg, G., Schmitz, M., 2012. The Geologic Time Scale 2012, 2-volume set. Elsevier. Graham, S.A., Hendrix, M.S., Johnson, C.L., Badamgarav, D., Badarch, G., Amory, J., Porter, M., Barsbold, R., Webb, L.E., Hacker, B.R., 2001. Sedimentary record and tectonic implications of Mesozoic rifting in southeast Mongolia. Geol. Soc. Am. Bull. 113, 1560–1579. Hasegawa, H., Tada, R., Jiang, X., Suganuma, Y., Imsamut, S., Charusiri, P., Ichinnorov, N., Khand, Y., 2012. Drastic shrinking of the Hadley circulation during the midCretaceous Supergreenhouse. Clim. Past 8, 1323–1337. Henkes, G.A., Passey, B.H., Grossman, E.L., Shenton, B.J., Pérez-Huerta, A., Yancey, T.E., 2014. Temperature limits for preservation of primary calcite clumped isotope paleotemperatures. Geochim. Cosmochim. Acta 139, 362–382. Hough, B.G., Fan, M., Passey, B.H., 2014. Calibration of the clumped isotope geothermometer in soil carbonate in Wyoming and Nebraska, USA: implications for paleoelevation and paleoclimate reconstruction. Earth Planet. Sci. Lett. 391, 110–120. Hu, G., Hu, W., Cao, J., Yao, S., Xie, X., Li, Y., Liu, Y., Wang, X., 2012. Deciphering the Early Cretaceous transgression in coastal southeastern China: constraints based on petrography, paleontology and geochemistry. Palaeogeogr. Palaeoclimatol. Palaeoecol. 317–318, 182–195. Hu, J.F., Peng, P.A., Liu, M.Y., Xi, D.P., Song, J.Z., Wan, X.Q., Wang, C.S., 2015. Seawater incursion events in a Cretaceous paleo-lake revealed by specific marine biological markers. Sci. Rep. 5. Huntington, K.W., Eiler, J.M., Affek, H.P., Guo, W., Bonifacie, M., Yeung, L.Y., Thiagarajan, N., Passey, B., Tripati, A., Daëron, M., Came, R., 2009. Methods and limitations of ‘clumped’ CO2 isotope (47) analysis by gas-source isotope ratio mass spectrometry. J. Mass Spectrom. 44, 1318–1329. Huntington, K.W., Lechler, A.R., 2015. Carbonate clumped isotope thermometry in continental tectonics. Tectonophysics 647–648, 1–20. Huntington, K.W., Wernicke, B.P., Eiler, J.M., 2010. Influence of climate change and uplift on Colorado Plateau paleotemperatures from carbonate clumped isotope thermometry. Tectonics 29, TC3005. Isozaki, Y., Aoki, K., Nakama, T., Yanai, S., 2010. New insight into a subductionrelated orogen: a reappraisal of the geotectonic framework and evolution of the Japanese Islands. Gondwana Res. 18, 82–105.

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Kim, J.H., van der Meer, J., Schouten, S., Helmke, P., Willmott, V., Sangiorgi, F., Koç, N., Hopmans, E.C., Damsté, J.S.S., 2010. New indices and calibrations derived from the distribution of crenarchaeal isoprenoid tetraether lipids: implications for past sea surface temperature reconstructions. Geochim. Cosmochim. Acta 74, 4639–4654. Kim, S.T., O’Neil, J.R., 1997. Equilibrium and nonequilibrium oxygen isotope effects in synthetic carbonates. Geochim. Cosmochim. Acta 61, 3461–3475. Lee, Y.S., Han, H.-C., Hwang, J.H., Kee, W.-S., Kim, B.C., 2011. Evidence for significant clockwise rotations of the Korean Peninsula during Cretaceous. Gondwana Res. 20, 904–918. Li, J., Zhang, Y., Dong, S., Johnston, S.T., 2014. Cretaceous tectonic evolution of South China: a preliminary synthesis. Earth-Sci. Rev. 134, 98–136. Linnert, C., Robinson, S.A., Lees, J.A., Bown, P.R., Pérez-Rodríguez, I., Petrizzo, M.R., Falzoni, F., Littler, K., Arz, J.A., Russell, E.E., 2014. Evidence for global cooling in the Late Cretaceous. Nat. Commun. 5. Meyer, H.W., 1992. Lapse rates and other variables applied to estimating paleoaltitudes from fossil floras. Palaeogeogr. Palaeoclimatol. Palaeoecol. 99, 71–99. Passey, B.H., Henkes, G.A., 2012. Carbonate clumped isotope bond reordering and geospeedometry. Earth Planet. Sci. Lett. 351–352, 223–236. Passey, B.H., Levin, N.E., Cerling, T.E., Brown, F.H., Eiler, J.M., 2010. High-temperature environments of human evolution in East Africa based on bond ordering in paleosol carbonates. Proc. Natl. Acad. Sci. 107, 11245–11249. Peters, N.A., Huntington, K.W., Hoke, G.D., 2013. Hot or not? Impact of seasonally variable soil carbonate formation on paleotemperature and O-isotope records from clumped isotope thermometry. Earth Planet. Sci. Lett. 361, 208–218. Pucéat, E., Lécuyer, C., Donnadieu, Y., Naveau, P., Cappetta, H., Ramstein, G., Huber, B.T., Kriwet, J., 2007. Fish tooth δ 18 O revising Late Cretaceous meridional upper ocean water temperature gradients. Geology 35, 107–110. Quade, J., Eiler, J., Daëron, M., Achyuthan, H., 2013. The clumped isotope geothermometer in soil and paleosol carbonate. Geochim. Cosmochim. Acta 105, 92–107. Quan, C., Liu, Z., Utescher, T., Jin, J., Shu, J., Li, Y., Liu, Y.-S., 2014. Revisiting the Paleogene climate pattern of East Asia: a synthetic review. Earth-Sci. Rev. 139, 213–230. Ringham, M.C., Hoke, G.D., Huntington, K.W., Aranibar, J.N., 2016. Influence of vegetation type and site-to-site variability on soil carbonate clumped isotope records, Andean piedmont of Central Argentina (32–34◦ S). Earth Planet. Sci. Lett. 440, 1–11. Snell, K.E., Koch, P.L., Druschke, P., Foreman, B.Z., Eiler, J.M., 2014. High elevation of the ‘Nevadaplano’ during the Late Cretaceous. Earth Planet. Sci. Lett. 386, 52–63. Song, Y., Stepashko, A.A., Ren, J., 2015. The Cretaceous climax of compression in Eastern Asia: age 87–89 Ma (late Turonian/Coniacian), Pacific cause, continental consequences. Cretac. Res. 55, 262–284. Tobin, T.S., Wilson, G.P., Eiler, J.M., Hartman, J.H., 2014. Environmental change across a terrestrial Cretaceous–Paleogene boundary section in eastern Montana, USA, constrained by carbonate clumped isotope paleothermometry. Geology 42, 351–354. von Huene, R., Ranero, C.R., 2009. Neogene collision and deformation of convergent margins along the backbone of the Americas. Geol. Soc. Am. Mem. 204, 67–83. Wang, C., Scott, R.W., Wan, X., Graham, S.A., Huang, Y., Wang, P., Wu, H., Dean, W.E., Zhang, L., 2013. Late Cretaceous climate changes recorded in Eastern Asian lacustrine deposits and North American Epieric sea strata. Earth-Sci. Rev. 126, 275–299. Wolfe, J.A., 1992. An analysis of present-day terrestrial lapse rates in the western conterminous United States and their significance to paleoaltitudinal estimates. Geol. Surv. Bull. 1964, 35. Wu, H., Zhang, S., Hinnov, L.A., Jiang, G., Yang, T., Li, H., Wan, X., Wang, C., 2014. Cyclostratigraphy and orbital tuning of the terrestrial upper Santonian–Lower Danian in Songliao Basin, northeastern China. Earth Planet. Sci. Lett. 407, 82–95. Xu, W.L., Pei, F.P., Wang, F., Meng, E., Ji, W.Q., Yang, D.B., Wang, W., 2013. Spatialtemporal relationships of Mesozoic volcanic rocks in NE China: constraints on tectonic overprinting and transformations between multiple tectonic regimes. J. Asian Earth Sci. 74, 167–193. Yang, Y.T., 2013. An unrecognized major collision of the Okhotomorsk Block with East Asia during the Late Cretaceous, constraints on the plate reorganization of the Northwest Pacific. Earth-Sci. Rev. 126, 96–115. Zhang, Y., Dong, S., Shi, W., 2003. Cretaceous deformation history of the middle Tan-Lu fault zone in Shandong Province, eastern China. Tectonophysics 363, 243–258. Zhang, J., Zhang, H.F., Ying, J.F., Tang, Y.J., Niu, L.F., 2008. Contribution of subducted Pacific slab to Late Cretaceous mafic magmatism in Qingdao region, China: a petrological record. Isl. Arc 17, 231–241. Zhang, S.H., Zhao, Y., Davis, G.A., Ye, H., Wu, F., 2014. Temporal and spatial variations of Mesozoic magmatism and deformation in the North China Craton: implications for lithospheric thinning and decratonization. Earth-Sci. Rev. 131, 49–87. Zhang, Y., Li, J., Zhang, T., Dong, S., Yuan, J., 2008. Cretaceous to Paleocene tectonosedimentary evolution of the Jiaolai Basin and the contiguous areas of the Shandong Peninsula (North China) and its geodynamic implications. Acta Geol. Sin. 82, 1229–1257 (in Chinese with English abstract).

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Zhu, G., Chen, Y., Jiang, D., Lin, S., 2015. Rapid change from compression to extension in the North China Craton during the Early Cretaceous: evidence from the Yunmengshan metamorphic core complex. Tectonophysics 656, 91–110.

Zhu, G., Jiang, D., Zhang, B., Chen, Y., 2012. Destruction of the eastern North China Craton in a backarc setting: evidence from crustal deformation kinematics. Gondwana Res. 22, 86–103.