An Early Pleistocene atmospheric CO2 record based on pedogenic carbonate from the Chinese loess deposits

Earth and Planetary Science Letters 426 (2015) 69–75

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Earth and Planetary Science Letters www.elsevier.com/locate/epsl

An Early Pleistocene atmospheric CO2 record based on pedogenic carbonate from the Chinese loess deposits Jiawei Da a , Yi Ge Zhang b,c , Hongtao Wang a , William Balsam d , Junfeng Ji a,∗ a

Key Laboratory of Surficial Geochemistry, Ministry of Education; School of Earth Sciences and Engineering, Nanjing University, Nanjing 210093, China Department of Geology and Geophysics, Yale University, New Haven, CT 06511, USA c Department of Earth and Planetary Sciences, Harvard University, Cambridge, MA 02138, USA d Department of Earth Sciences, Dartmouth College, Hanover, NH, USA b

a r t i c l e

i n f o

Article history: Received 16 October 2014 Received in revised form 28 May 2015 Accepted 29 May 2015 Available online xxxx Editor: G.M. Henderson Keywords: atmospheric carbon dioxide pedogenic carbonate Chinese Loess Plateau Early Pleistocene

a b s t r a c t The tight coupling between temperature and atmospheric CO2 is shown by ice core records for the past 0.8 million years (Myr). However, the modern atmospheric partial pressure of CO2 (pCO2 ) has exceeded previous interglacial pCO2 levels over the past 0.8 Myr, suggesting that the earlier part of the Pleistocene and Pliocene might be a better analog of today’s radiative forcing of CO2 . The early Pleistocene experienced a constant cooling characterized by the intensification of northern hemisphere glaciations. Existing pCO2 records developed from marine organic matter and inorganic precipitates, however, disagree with the trends and absolute values of CO2 over this time interval. Here we present quantitative interglacial pCO2 estimates from ∼2.6–0.9 Myr using the stable carbon isotopic compositions of pedogenic carbonates collected from the Chinese Loess Plateau (CLP). Our pCO2 records provide the first documentation of pCO2 from continental sedimentary deposits over the early Pleistocene. The successive decrease of our pCO2 records is broadly consistent with the increase in deep-sea δ 18 O and the overall decline of sea surface temperature (SST) at this time, but in contrast with the increasing peak interglacial pCO2 recorded in ice cores for the last 0.8 Myr. © 2015 Elsevier B.V. All rights reserved.

1. Introduction Atmospheric CO2 levels over the past 0.8 million years (Myr) achieved by ice cores are in good agreement with Antarctic temperature shown by δ D of ice (Lüthi et al., 2008), indicating a strong coupling between pCO2 and climate response. Beyond the 0.8-million-year history available from the ice cores, the pCO2 history remains vague and largely relies on geochemical or paleontological proxies. Since modern pCO2 is substantially higher than peak interglacial pCO2 values recorded by the ice cores (Solomon et al., 2007), it is crucial to obtain better constraints on pCO2 in the geological past, especially during the early Pleistocene and Pliocene epochs that have higher CO2 levels than the late Pleistocene but share similar boundary conditions (e.g., plate configurations) with the modern Earth. Previous reconstructions determined early Pleistocene pCO2 mainly through two proxies, the boron isotope composition of planktonic foraminifera (Hönisch et al., 2009; Seki et al., 2010; Bartoli et al., 2011; Martínez-Botí et al., 2015) and the carbon iso-

*

Corresponding author. E-mail address: [email protected] (J. Ji).

http://dx.doi.org/10.1016/j.epsl.2015.05.053 0012-821X/© 2015 Elsevier B.V. All rights reserved.

tope composition of alkenones (Pagani et al., 2010; Seki et al., 2010; Zhang et al., 2013; Badger et al., 2013), both of which are applied to marine sediments. However, uncertainties associated with the two proxies remain large. Both proxies are based on assumptions which could be invalidated by local or regional conditions, e.g., air–sea equilibrium is difficult to achieve due to the biological pump and/or upwelling of CO2 -rich deep waters (Takahashi et al., 2009). For alkenones, physiological conditions of the alkenone-producing haptophyte algae are difficult to constrain, e.g., growth rate, level of irradiance (Pagani et al., 2014). For boron, in addition to carbonate dissolution and diagenesis of foraminiferal tests, secular changes in seawater δ 11 B are still unclear and impose a large uncertainty to pCO2 estimates (e.g. Pearson et al., 2009). Available boron- and alkenone pCO2 reconstructions display disparate pCO2 trends during the early Pleistocene with the boron proxy indicating a sudden decline of pCO2 at ∼2.1 Myr (Bartoli et al., 2011), whereas the alkenone proxy indicates a gradual decline between ∼2.5–1.5 Myr (Pagani et al., 2010; Zhang et al., 2013). Therefore, the trends and magnitude of pCO2 changes remain ambiguous during the early Pleistocene. Hence, a top priority is to use another method that provides an independent constraint on the evolution of pCO2 .

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The pedogenic carbonate paleobarometer has been widely used to reconstruct ancient atmospheric CO2 concentrations (e.g. Cerling, 1992; Ekart et al., 1999; Nordt et al., 2002; Retallack, 2009). In order to calculate pCO2 , δ 13 C values of inorganic carbonate and soil organic matter need to be measured. However, due to the uncertainty of S ( z) (soil-respired CO2 ), one of the parameters in the barometer equation which cannot be directly measured from the paleosols, pedogenic carbonate-derived pCO2 reconstruction usually carries larger uncertainties compared to other proxies. Recent investigations have made substantial progress on the quantification of S ( z), such as using mean annual precipitation (MAP) as an alternate proxy to calculate S ( z) (Cotton and Sheldon, 2012), and constrain soil-specific S ( z) with modern soil samples (Montañez, 2013). These combined efforts have greatly improved the pedogenic carbonate proxy towards resolving the long-lasting discrepancy of pCO2 estimates between carbonate-pCO2 and other proxies. Here we present pCO2 reconstructions from pedogenic carbonates collected on the Chinese Loess Plateau (CLP), using the paleosol paleobarometer. Carbonate nodules are widely dispersed in almost every loess-paleosol strata in the two sections we studied, enabling the construction of two pCO2 curves from ∼2.6–0.9 Myr, a time when the global temperature experienced a unidirectional cooling after the intensification of northern hemisphere glaciations. The purpose of this work is twofold: (1) to test the sensitivity of soil carbonate to relatively small changes in atmospheric CO2 levels in the Pleistocene; (2) to provide the unique opportunity of critically evaluating the role of pCO2 variations in the early Pleistocene global cooling.

Fig. 1. Map showing the Chinese Loess Plateau (CLP) with the location of the Baoji and Lingtai section, and the Pacific Ocean with the location of Ocean Drilling Program (ODP) sites 1143, 1012 and 846 (Li et al., 2011; Brierley and Fedorov, 2010; Lawrence et al., 2006).

2. Samples and methods

when precipitation intensifies, primary carbonates are dissolved in the soil water, along with CO2 released by soil respiration. The gravity-driven soil water percolates downward, and eventually reprecipitates as pedogenic carbonate when evapotranspiration exceeds precipitation. Due to carbon isotope (13 C/12 C) equilibrium between pedogenic carbonate and soil CO2 during the time of crystallization, and the fact that soil CO2 is a mixture of atmospheric CO2 and soil-respired CO2 , we are capable of calculating pCO2 with measured carbon isotope composition of pedogenic carbonate, using a diffusion–reaction model developed by Cerling (1991):

2.1. Chinese loess and the East Asian monsoon

[CO2 ]atm = S ( z)

In northern China, Quaternary loess-paleosol sequences (S— interglacial paleosol unit, L—glacial loess unit) are well known for their great thickness, high resolution and richness of paleoenvironmental information. Large amount of investigations based on these units have been published concerning the history and variability of the East Asian monsoon (e.g., Liu, 1985; Liu and Ding, 1998; Ji et al., 2001; Guo et al., 2002). The loess horizons, characterized by yellowish color, massive structure and high carbonate content, are interpreted to have accumulated during dry, cold glacial periods with a significant weakening of Asian summer monsoon, whereas the paleosol horizons, characterized by brownish or reddish color and pedogenic structures, formed during interglacial periods when summer monsoon were intensified (Liu, 1985). The East Asian summer monsoon brings warm, moisture-laden air masses from tropical oceans to the CLP, resulting in substantial precipitation, while the winter monsoon winds flowing from the Siberian region prevail on the CLP, leading to a dry and cold climate. This monsoonal climate, characterized by large seasonal variations of precipitation, facilitates pedogenic carbonate precipitation on the CLP (Liu, 1985). The size and illuviation depth of pedogenic carbonates in the CLP varies with MAP, with the largest and deepest found in the Southeastern part of the CLP because of the relatively warm and humid climate (Zhao, 1992). In some well-developed paleosols, the carbonate is depleted in the eluvial horizon and reprecipitated carbonates appear as calcium carbonate concretions near the bottom of the B horizon (Han et al., 1997).

where S ( z) is the concentration of CO2 derived from soil respiration (S ( z) = [CO2 ]soil − [CO2 ]atm ) at depth z; δ 13 C is the carbon isotope composition of the total soil CO2 (s), soil-respired CO2 (r), and atmospheric CO2 (a). The 1.0044 is the mixing ratio of diffusion coefficients of 12 CO2 to 13 CO2 and 4.4 is the fractionation associated with different diffusivity between them. δ 13 Cs is calculated from the measured carbon isotopic composition of pedogenic carbonate (δ 13 Cc ), using a temperature dependent carbon fractionation factor (Romanek et al., 1992). δ 13 Cr is determined by the carbon isotopic composition of soil organic matter (δ 13 Csom ) occluded in the carbonate nodules (Breecker, 2013).

2.2. Pedogenic carbonate pCO2 paleobarometer Pedogenic carbonate is secondary carbonate formed in calcareous soils, or noncalcareous soils where Ca is provided from external sources such as precipitation or dust. During the growth season

δ 13 Cs − 1.0044δ 13 Cr − 4.4 δ 13 Ca − δ 13 Cs

2.3. Samples To reconstruct atmospheric CO2 concentration in the early Pleistocene, carbonate nodules were sampled from the loesspaleosol sections in Lingtai (35◦ 04 N, 107◦ 39 E, 1350 m above sea level) and Baoji (34◦ 22 N, 107◦ 14 E, 850 m above sea level) (Fig. 1); the sections have thicknesses of about 176 m and 161 m respectively. The two sections are located in southern part of the CLP and the paleosols are characterized by a well-developed Bt horizon with continuous clay coatings on ped faces, indicating a forest or steppe-forest ecosystem under a warm and humid climate, conditions that are suitable for soil carbonate-pCO2 approach. We collected samples of carbonate nodules from the paleosolloess sequences from S1 through L33. They are typical CLP pedogenic carbonates and contain little detrital carbonates (Zheng et al., 1987) which might have a very different carbon isotopic composition compared to pedogenic carbonates. There are two types of pedogenic carbonates in the two sections. One is semi-round or ginger-shaped carbonate nodules ranging from a few centimeters

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to tens of centimeters in diameter and are commonly scattered at the bottom of the respective paleosol layer. The other type is continuous carbonate nodules forming a horizontal carbonate horizon with a thickness of 20–80 cm. It is noteworthy that, although some of the carbonate nodule horizons are located in the underlying loess units due to long leaching depth, they were thought to be formed by strong leaching and illuviation during the development of the paleosols (Zhao, 1992). This is likely the case in S1, S3 and S4 in Baoji sections where we found carbonate nodule samples. Therefore, we assume that all of our carbonate samples were formed during interglacial periods throughout the Pleistocene. It is possible that during glacial periods, occasionally elevated precipitation and favorable temperature could facilitate nodule formation (Rowe and Maher, 2000). However, carbonate nodules formed under these conditions are probably much smaller (a few millimeters in size), different from the nodules that we collected for stable isotope analysis. This assumption is consistent with the modern observation that in the northwestern part of the CLP, where MAP is 200–400 cm, small carbonate nodules were found in the modern soil (Liu, 1985). We therefore conclude that the studied soil carbonate nodules were primarily formed during interglacial periods. Two or three samples were taken from each carbonate layer for replicate analysis. Nearly all the carbonate nodules we collected from the two loess-paleosol sections were formed at a depth greater than 50 cm in the soil profile, indicating a constant isotopic equilibrium according to CO2 diffusion models (Sheldon and Tabor, 2009). 2.4. Chronology The chronological framework of loess-paleosol sequences in the two sections are based on the paleomagnetic reversal sequences, which were detailed by previous studies (Rutter et al., 1991; Ding et al., 1999). The Brunhes–Matuyama reversal occurs at the base of L8 or the upper part of S8. The Jaramillo subchron lies between the upper part of L10 and the base of S11, while the Olduvai subchron is defined between the middle part of L25 and the lower part of S26. A basal age of ∼2.6 Myr is estimated for the Chinese loess deposits, based on the Matuyama/Gauss magnetic reversal occurring in the oldest loess unit, L33. The stratigraphy of loess-paleosol sequences is further delineated by magnetic susceptibility data; higher bulk magnetic susceptibility occurs in paleosols than in loess beds. Previous investigations of loess stratigraphy have revealed a good correlation between Chinese loess-paleosol sequences and global glacial cycles as recorded by δ 18 O records from deep sea sediments (Kukla and Cílek, 1996; Rutter et al., 1991), supporting the robustness of our chronology. The magnetic susceptibility data of our study sites provide a much more detailed age control for the Lingtai and Baoji section (Fig. 2). Spatial offset between carbonate dissolution and reprecipitation causes temporal offset between the formation age of carbonate nodule and the strata where it occurs. The absolute formation time of the carbonates is therefore difficult to assess. Recent investigations (e.g. Gocke et al., 2012) suggest that the age determination of pedogenic carbonates has an uncertainty of at least a millennia. Here we use the time range of the paleosol unit to represent the formation time of the carbonate nodules occurred within. For the convenience of comparing the reconstructed pCO2 between the two sections and with other marine derived pCO2 records, we give carbonate nodules from each paleosol unit a specific age rather than a time range, using the mean age of each paleosol unit’s time span (Fig. 2).

Fig. 2. Comparisons of our pedogenic carbonate-derived pCO2 record from Lingtai (upper) and Baoji sections (down) with benthic δ 18 O (Lisiecki and Raymo, 2005) and magnetic susceptibility. Horizontal grey lines are preindustrial pCO2 and δ 18 O level.

2.5. Carbon isotope analysis To obtain δ 13 Cr and δ 13 Cs for pCO2 calculation, we measured the δ 13 C values of pedogenic carbonates and the organic matter occluded in the carbonate nodules. For carbonate isotopic analyses, dried samples were homogenized by crushing and grinding into powder. ∼15 mg samples were analyzed on a ThermoFinnigan MAT 253 isotope ratio mass spectrometer using a Kiel IV carbonate device (75 ◦ C reaction in 100% H3 PO4 ). For organic carbon isotopic analysis, the bulk samples were dissolved in 7% HCl and ultrasonically rinsed with deionized water to remove carbonates (Cotton and Sheldon, 2012). The dried samples were then homogenized by grinding and loaded into tin capsules and analyzed on a Thermo Delta V+ isotope ratio mass spectrometer, using a Costech elemental analyzer attached to it. Both carbonate and organic carbon isotopic results are reported in permil (h) notation relative to the Pee Dee Belemnite (VPDB)

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Table 1 Values of S ( z) from carbonate nodules collected from loess-paleosol units S1–S7 in the CLP and calculated with the inverted paleobarometer equation. Localities

Loess-paleosol unit

Sample

Age (Myr)

δ 13 Cc

δ 18 Oc

Lingtai

S1

1 2 3 4 5 6 7 8 9 10 11 12 13

0.12 0.12 0.12 0.33 0.33 0.42 0.42 0.78 0.78 0.12 0.24 0.33 0.42

−7.98 −7.87 −7.65 −7.08 −7.11 −7.76 −7.79 −8.75 −8.80 −7.23 −6.73 −5.78 −8.55

−8.77 −9.07 −8.83 −8.70 −9.03 −9.38 −9.24 −9.25 −9.06 −9.29 −9.04 −9.52 −9.43

S3 S4 S7 Baoji

S1 S2 S3 S4

δ 13 Csom −23.83 −23.71 −23.71 −23.55 −23.88 −24.43 −23.58 −24.20 −23.94 −24.09 −22.36 −22.46 −24.58

13 C

pCO2 (ppm)

δ 13 Cs

(21 ◦ C )

S ( z) (ppm)

15.85 15.85 16.07 16.47 16.77 16.68 15.78 15.46 15.14 16.85 15.63 16.67 16.03

270 270 270 266 266 274 274 248 248 270 252 266 274

−17.44 −17.33 −17.11 −16.54 −16.57 −17.22 −17.25 −18.21 −18.26 −16.69 −16.19 −15.24 −18.01

1413 1398 1239 985 889 1004 1453 1703 2103 888 1305 798 1382

Note: pCO2 records are from ice core data compiled in Lüthi (2008). δ 13 C values are reported in h vs PDB.

standard with a precision better than ±0.2h from the duplicate analyses. 2.6. pCO2 calculation The uncertainty associated with S ( z) estimates has been recognized as the primary contributor to the pCO2 reconstructions using the paleosol-pCO2 method, and in previously work, it is either assumed to be a certain number (e.g. Ekart et al., 1999), or calculated with an alternate proxy such as MAP (Cotton and Sheldon, 2012). Here we adopt a different approach of calibrating the S ( z) in Lingtai and Baoji sections by back calculating S ( z) using the isotope data and the known pCO2 values from ice cores (Montañez, 2013). Specifically, paleosol-loess sequences S0–L8 span the last 0.8 Myr, overlapping with the ice core records. Therefore, with measured carbon isotopic composition of samples from S0–L8, we are able to back-calculate S ( z) through the inverted barometer equation described above. We used the average pCO2 ice core value of each interglacial, together with the δ 13 Cs and δ 13 Cr values of our samples from each corresponding interglacial to calculate S ( z). The averaged S ( z) value for the past 0.8 Myr was then used to calculate pCO2 in the early Pleistocene, with the range of possible S ( z) providing the error envelope associated with the CO2 estimates. The average pCO2 uncertainty associated with different S ( z) is 140 ppmv. Detailed results of our S ( z) estimates are shown in Section 3.2. The isotopic equilibrium between soil CO2 and carbonates is also temperature dependent (Romanek et al., 1992). Previous studies provide some helpful information on the temperature variation of soil carbonate formation on the CLP. Suarez et al. (2011) used the carbonate clumped isotope thermometer, with pedogenic carbonates collected in Lantian and Baode on the CLP, to reconstruct formation temperatures of samples found between 7–3 Myr. The resulting variations of temperatures from both localities are within 2 ◦ C, suggesting small changes of nodule formation temperature over Miocene–Pliocene. Furthermore, the clumped-isotope temperatures indicate that pedogenic carbonates are formed during warm and dry periods in the CLP. Therefore, we use the modern September and October mean air temperature (21 ◦ C) in Xi’an (34◦ 16 N, 108◦ 57 E), which is located near the two studied sites with good meteorological data services, as the formation temperature of pedogenic carbonates. To examine the sensitivity of pCO2 estimates related to the uncertainty of temperature, we found that a ±2 ◦ C temperature change would change our pCO2 estimates approximately ±35 ppm. Hence, the uncertainty of pCO2 estimates associated with soil temperature is minor compared to the uncertainty related to the S ( z) values. δ 13 Cr is usually approximated from the carbon isotope composition of soil organic matter (Breecker, 2013). We used δ 13 CSOM

occluded in the carbonate nodules. δ 13 C of atmospheric CO2 exhibits minimal changes during the time interval we focus on (e.g., Tipple et al., 2010), and δ 13 Ca proved to be one of the least sensitive variables in the equation (Cerling, 1992). We therefore used the preindustrial value of −6.5h. 3. Results With the stable isotope measurements and other necessary information we were able to calculate pCO2 for all 120 samples, covering the time period of ∼0.9–2.6 Myr. However, not all samples are suitable for CO2 estimates due to the complexity of soil carbonate formation. Cerling (1991) examined the isotopic fractionation between soil-respired CO2 and soil carbonate in isotopic equilibrium (13 C), at soil temperatures 0–30 ◦ C. Pedogenic carbonates, with 13 C values (δ 13 Cc − δ 13 Corg ) outside the range between 14h and 17h, are considered to have formed beyond an equilibrium state and therefore not applicable for the diffusion–reaction model. However, carbonates with 13 C values >17h are likely to precipitate under a low soil productivity system where the isotopically heavier atmospheric CO2 occupied a considerable proportion of soil CO2 (Sheldon and Tabor, 2009), and thus is still valid for the paleobarometer approach. Alternatively, it is also possible that these samples were formed beyond a two-component system and are inappropriate for pCO2 reconstruction. Therefore, despite their potential utility, we abandoned these samples. Altogether, 37 samples were screened out by this criterion which accounts for 31% of the samples being tested. 3.1. The stable isotopic compositions of pedogenic carbonates Pedogenic carbonate samples from two sections were divided into two groups, one collected from S0–L8 used for S ( z) estimates and the other from S9–L33 used for pCO2 reconstructions. As shown in Table 1 and Table S1, δ 13 Cc values of the former group range between −8.8h and −5.78h (PDB), with an average of −7.62h, whereas δ 13 Cc values of the latter group range between −9.6h and −6.14h and averages −8.2h. On the other hand, the δ 13 Csom values of the former group range between −23.83h and −24.58h with an average of −23.72h, while the δ 13 Csom values of the latter group range between −22.87h and −25.33h and averages −24.27h. For oxygen isotopic compositions, the former group gives a range between −9.52h and −8.7h, and the latter group ranges between −9.58h and −7.69h, with averages of −9.12h and −8.98h respectively.

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3.2. S ( z) estimates and pCO2 reconstruction Between 0.8 and 0 Ma, the S ( z) ranges from 798 to 2103 ppmv (Table 1), with an average value of 1274 ppmv, with the lowest value occurred in S2 from the Baoji section, and the highest in S7 at Lingtai section. Calculated S ( z) values between the two sections are comparable with each other. The average value and the range of S ( z) were used to reconstruct early Pleistocene pCO2 values from both sections (Fig. 2). In Lingtai section, pCO2 remained at a relatively high level of 400 ppm before 2.2 Myr. However, pCO2 decreased to a preindustrial level of ∼280 ppm between 2.2–1.6 Myr. After 1.6 Myr, pCO2 continued to decrease, reaching a minimum value of ∼150 ppm at 1.3 Myr. The pCO2 trend reconstructed from Baoji section is similar to Lingtai section, indicating a successive decrease during the early Pleistocene. The absolute values of pCO2 during 1.5–1.3 Ma are unreasonably low compared to ice core records. These low CO2 estimates probably result from the large uncertainty of S ( z) values, which yields an average error of ±140 ppm. The broad agreement between pCO2 records from Baoji and Lingtai sections supports the robustness of our methodology. These CO2 records show a long-term decrease of interglacial pCO2 , from as high as 400 ppm to glacial–interglacial CO2 levels in the early Pleistocene (Fig. 2). 4. Discussion 4.1. Examining the fidelity of S ( z) estimates S ( z) is dependent on many factors such as biological productivity, soil porosity, and tortuosity (Cerling, 1991). Controls on S ( z) variations are comparable with controls on soil productivity, such as temperature, precipitation and local ecosystems (Leith and Whittaker, 1975). Recent studies of modern soils (Breecker et al., 2009; Montañez, 2013) have revealed that pedogenic carbonates are not formed during the mean growing season, but during warm, dry periods when precipitation decreases along with the soil’s respiration rate. Therefore, the S ( z) value during the time of carbonate precipitation is much lower than previously assumed (Cerling, 1991; Ekart et al., 1999). The estimated S ( z) values at Lingtai and Baoji sections range from 798 to 2103 ppm with an average of 1271 ppm. This level of S ( z) is consistent with recent paleo-studies (Montañez et al., 2007; Schaller et al., 2011) and observations of modern soils (Breecker et al., 2009; Montañez, 2013). Specifically, our S ( z) result falls inside the S ( z) range of 500–2500 ppm, suggested by Montañez (2013) as the best estimate for Calcisols, a paleosol-type closely resembles the paleosols in this study. Relatively low S ( z) values are further supported by in-situ measurements in Chinese loess near Beijing in the north China which yielded a range of 1434–2458 ppm (Liu et al., 2000). These measurements were performed during September at an average depth of 1 m. It is known that the concentration of soil CO2 becomes relatively constant below 50 cm (Cerling, 1992), therefore the measured S ( z) matched the formation depth of our carbonate samples, which are found 50 cm below the top of the paleosol profile. Although the loess-paleosol sequence in Beijing does not necessarily reflect conditions on the CLP, the two regions share similarities between the soil type and climatic conditions during which soil carbonate precipitates, therefore S ( z) values should be comparable. In addition, several lines of evidence suggest that our S ( z) estimates derived from the past 0.8 Myr are also applicable to the time interval 2.6–0.9 Myr. Our results show that the organic carbon isotopic compositions between the last 0.8 Myr and 0.9–2.6 Myr at both sections are consistent with each other, indicating a stable

Fig. 3. Multiproxy Pleistocene pCO2 evolution. The grey curve is the ice core pCO2 record (Lüthi et al., 2008). pCO2 results from Baoji and Lingtai are shown separately using the averaged value for each time interval. Blue and red icons represent glacial and interglacial pCO2 values, respectively. Our data is compared to the (middle) alkenone-derived pCO2 records (Site 925a from Zhang et al., 2013; Sites 925b, 806 and 1202 from Pagani et al., 2010) and (upper) boron-derived pCO2 records (Hönisch et al., 2009; Seki et al., 2010; Bartoli et al., 2011). Note the different scales for each panel. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

C3-dominated ecosystem throughout the Pleistocene. Furthermore, δ 18 Oc is well correlated with the isotopic composition of meteoric water and could be used to reconstruct paleotemperature worldwide (Dworkin et al., 2005). Our δ 18 Oc results of the last 0.8 Myr are similar to those from 2.6–0.9 Myr (Table 1 and Table S1), suggesting that changes in the formation temperature of carbonate nodules were minor. Given constant soil types and steady environmental conditions indicated by stable isotope results, it is reasonable to use S ( z) estimates of the last 0.8 Myr for pCO2 reconstruction during the early Pleistocene. 4.2. Comparisons with published pCO2 records We compare our pCO2 reconstruction based on terrestrial carbonate nodules from central China to published pCO2 results, which were derived from marine sediments utilizing either alkenone or boron approaches (Hönisch et al., 2009; Seki et al., 2010; Bartoli et al., 2011; Pagani et al., 2010; Zhang et al., 2013). All records show a distinct decrease of CO2 since 2.2 Myr (Fig. 3). However, the duration and magnitude of this drop vary among different approaches. The alkenone-based pCO2 record represents a mild and successive decreasing trend from 400 ppm to 300 ppm between 2.6 and 1.5 Myr. This continuous CO2 drop agrees with our carbonate-based results. Between 1.5 and 1.2 Myr, alkenone-CO2 records show a brief increase, then declined again towards the late Pleistocene. This increase is also seen in our results, but the timing is slightly older in our data, about 1.6 Myr. Unlike alkenone-derived pCO2 data, the boron isotope record shows an abrupt decline to interglacial pCO2 at ∼2.1 Myr, from values as high as ∼400 ppm to a progressively decreasing Pleistocene value of 250 ppm, coinciding with the most rapid decrease of CO2 in our carbonate-derived results. However, boron-based pCO2 remained steady between ∼250–300 ppm since 2.0 Myr,

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Fig. 4. pCO2 in this study compared to marine SSTs from ODP sites 1143, 846 and 1012 (Li et al., 2011; Brierley and Fedorov, 2010; Lawrence et al., 2006; Fedorov et al., 2013). pCO2 results from Baoji and Lingtai are shown separately, using an averaged value for each time interval. The grey curve is the ice core pCO2 record (Lüthi et al., 2008). Vertical grey bar marks the time period (2.2–2.0 Myr) of substantial cooling indicated from SSTs and pCO2 records.

whereas our result continued to drop, similar to the alkenonederived results. 4.3. Comparisons between pCO2 , benthic δ temperature records

18

O and sea surface

The long-term cooling trend in the Cenozoic is closely coupled with the decrease of atmospheric CO2 , from as high as >1000 ppm in the Eocene to the preindustrial level of 280 ppm (Pagani et al., 2005). Growing evidence has confirmed the critical role of atmospheric CO2 on the onset of major glaciations. DeConto et al. (2008) used an ice-sheet model to show that the long-term decrease of Cenozoic pCO2 levels was responsible for the growth of ice sheets, with ∼750 and ∼280 ppm as the threshold of major Antarctic and Northern Hemisphere glaciations (NHG). Lunt et al. (2008) also suggested that the decrease of pCO2 from mid-Pliocene to Quaternary triggered the Greenland glaciation. However, the relationship between global temperature, ice sheet variability and CO2 after the onset of major NHG is still unclear. Our CO2 records provide the opportunity to investigate the relationship between pCO2 , benthic δ 18 O and sea surface temperature (SST) in the early Pleistocene. Benthic δ 18 O stack (Lisiecki and Raymo, 2005) shows a successive increasing trend from 2.6–1.5 Myr (Fig. 2), indicating high latitude cooling and/or the expansion of ice sheet volume, including both glacial and interglacial periods during the early Pleistocene. This overall trend is in agreement with our pCO2 result (Fig. 4), suggesting a coupling between pCO2 and the global cooling trend since the onset of NHG at ∼2.7 Myr. However, recent studies suggest that sea level and deep-sea temperature, the two major components that drive the variations of benthic δ 18 O (e.g. Rohling et al., 2014), do not always track each other. Rohling et al. (2014) argued that the marked positive shift of benthic δ 18 O ∼2.73 Myr ago was caused by deep sea

cooling rather than glaciation process. Instead, the first major NHG occurred at 2.15 Myr ago. If valid, this glaciation coincided with a major step of pCO2 drawdown from 2.2–2.0 Myr as shown by our results, as well as the boron-derived records (Bartoli et al., 2011). In addition to temperature estimates from the benthic δ 18 O records, ocean temperatures and temperature distributions can also be reconstructed directly using SST proxies applied to marine sediments. We selected Pleistocene SST records from key regions of the low-latitude Pacific Ocean with adequate temporal resolution: Site 1143 from the South China Sea is within the West Equatorial Pacific warm pool (Li et al., 2011), Site 846 from the East Equatorial Pacific coastal upwelling zone (Lawrence et al., 2006), and Site 1012 from the subtropical upwelling zone in the northeast Pacific (Brierley and Fedorov, 2010). All the SSTs are derived from the alkenone unsaturation index. All time series have been interpolated to an even spacing of 2 kyr by Fedorov et al. (2013). As shown in Fig. 4, all SSTs showed slight decline from 2.6–1.6 Myr, in accordance with our pCO2 decrease, indicating a coupling between atmospheric CO2 concentration and low latitude temperature. The most rapid drawdown of 100 ppm pCO2 levels during 2.2 and 2.0 Myr coincides with the substantial cooling at sites 846 and 1012 from the Pacific upwelling regions. In general, the broad agreement between pCO2 , benthic δ 18 O and marine SST records suggest that atmospheric CO2 , as an important climate forcing, imposed significant impact on the highlatitude ice-sheet build up and low-latitude temperature in the early Pleistocene when the atmospheric CO2 levels were higher than the late Pleistocene glacial–interglacial cycles and more comparable with today’s condition. This result is consistent with our current understanding that atmospheric CO2 is an important climate forcing in the past and future. 5. Conclusion Our stable isotope study from the Chinese Loess Plateau demonstrates that the pedogenic carbonate pCO2 proxy can be applied to the Pleistocene where CO2 changes are relatively small. Our interglacial pCO2 result shows a unidirectional, stepwise decrease during the early Pleistocene, from as high as 400 ppm at the Pliocene–Pleistocene boundary, to as low as 150 ppm right before the onset of Mid-Pleistocene Transition. The long-term pCO2 decrease as indicated by our pedogenic carbonate proxy reconstruction is in broad agreement with alkenone- and boron-derived pCO2 records, but also show distinct features. An overall coupling between pCO2 , deep-sea δ 18 O and marine SST suggests that atmospheric CO2 forcing played a significant role in the early Pleistocene cooling. Acknowledgements This study was funded by the National Natural Science Foundation of China through grants 41273111, 41230526 and 41321062. Y.G.Z. thanks the Yale Graduate Fellowship and Harvard Center for the Environment Fellowship for support. We thank the Editor (Gideon Henderson) and three anonymous reviewers for their thoughtful comments. Appendix A. Supplementary material Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2015.05.053. References Badger, M.P.S., Schmidt, D.N., Mackensen, A., et al., 2013. High-resolution alkenone palaeobarometry indicates relatively stable pCO2 during the Pliocene (3.3–2.8 Ma). Philos. Trans. R. Soc., Math. Phys. Eng. Sci. 371 (2001), 20130094.

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