Journal of Geophysical Research: Solid Earth

RESEARCH ARTICLE An Integrated Study of the Eolian Dust in Pelagic Sediments 10.1002/2017JB014951 From the North Pacific Ocean Based on Environmental Key Points: Magnetism, Transmission Electron Microscopy, • Ferrimagnetic and antiferromagnetic minerals were identified and Diffuse Reflectance Spectroscopy systematically in North Pacific Ocean pelagic sediments Qiang Zhang1,2 , Qingsong Liu3,4 , Jinhua Li2,4, and Youbin Sun5 • Ferrimagnetic minerals record the overall trend of increased eolian dust 1College of Earth and Planetary Sciences, University of Chinese Academy of Sciences, Beijing, China, 2State Key Laboratory inputs since the late Pliocene 3 • A parameter obtained from diffuse of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, China, Department 4 reflectance spectroscopy is proposed of Ocean Science and Engineering, Southern University of Science and Technology, Shenzhen, China, Laboratory for to trace eolian dust in pelagic Marine Geology, Qingdao National Laboratory for Marine Science and Technology, Qingdao, China, 5State Key Laboratory sediments of Loess and Quaternary Geology, Institute of Earth Environment, Chinese Academy of Sciences, Xi’an, China

Supporting Information: fi • Supporting Information S1 Abstract Eolian dust is a major terrigenous component in North Paci c Ocean pelagic sediments and is an • Table S1 important recorder of Asian terrestrial environmental evolution and Northern Hemisphere atmospheric circulation. In order to extract and quantify wind-borne mineral signals and develop a reliable eolian dust Correspondence to: proxy for pelagic sediments, we investigated sediments from Ocean Drilling Program Hole 885A, North Q. Liu, fi [email protected] Paci c Ocean, by integrating results from environmental magnetism, transmission electron microscopy, and diffuse reflectance spectroscopy (DRS). Our results indicate that ferrimagnetic and antiferromagnetic minerals coexist in the sediments. The former includes biogenic magnetite, nano-sized single domain (titano) Citation: fi Zhang, Q., Liu, Q., Li, J., & Sun, Y. (2018). magnetite inclusions embedded in silicate hosts, submicron vortex state magnetite, and ultra ne An integrated study of the eolian dust in superparamagnetic maghemite. The latter includes hematite and goethite. Our results indicate that magnetic pelagic sediments from the North properties can be used to determine an overall trend of increased eolian dust inputs, with additional fi Paci c Ocean based on environmental fi fl magnetism, transmission electron signi cant in uence from biogenic and volcanogenic components. In contrast, a newly proposed parameter fl microscopy, and diffuse re ectance RelHm+Gt (combined hematite and goethite concentration) obtained from DRS measurements extracts spectroscopy. Journal of Geophysical precisely the eolian signal and precludes other components. Our DRS results indicate that hematite/goethite – Research: Solid Earth, 123, 3358 3376. fi https://doi.org/10.1002/2017JB014951 and siliceous fractions deposited at Hole 885A were jointly delivered to the North Paci c Ocean from Asian dust sources. Therefore, we suggest that RelHm+Gt is a reliable eolian dust proxy for pelagic sediments in Received 7 SEP 2017 this region. Accepted 23 MAR 2018 Accepted article online 30 MAR 2018 Published online 29 MAY 2018 1. Introduction Eolian dust preserved in terrestrial stratigraphic profiles (An et al., 2001; Guo et al., 2002; Liu, 1985), ice cores (Biscaye et al., 1997; Bory et al., 2003), and marine sediments (Bailey et al., 2011; Larrasoaña et al., 2003; Rea, 1994; Roberts, Rohling, et al., 2011) carries considerable information about the aridification history of source areas and atmospheric circulation over different timescales. The concentration and grain size of eolian dust are linked strongly with environment and climate. For example, eolian dust fluxes in glacial stages tend to be greater than in interglacial periods (Lambert et al., 2008; Maher et al., 2010). In addition, the grain size of eolian dust can be used to semiquantitatively identify changes in atmospheric circulation intensity (Janecek & Rea, 1985; Sun et al., 2006). Moreover, eolian dust can influence climate and ocean ecosystems directly by changing the properties of the atmosphere (Spracklen et al., 2008) and by influencing the biogeochemistry of seawater (Jickells et al., 2005). Present-day global eolian dust sources are characteristic by a zonal distribution that extends from North Africa to eastern China (Prospero et al., 2002). As shown in satellite images, eolian dust masses can be trans- ported from Asian inland sources to the North Pacific Ocean, with activity greater in springtime (Husar et al., 1997; Prospero et al., 2002). The Taklimakan Desert, Gobi Desert, Hexi Corridor, Chinese Loess Plateau, and Mongolian Plateau are major representative East Asian dust sources (Prospero et al., 2002; Wang et al., 2004). Deep-sea basins with stable sedimentary environments in the North Pacific Ocean lie downwind from the major Asian dust sources (Tanaka & Chiba, 2006) and receive abundant wind-borne minerals that are fl fi ©2018. American Geophysical Union. deposited on the sea oor. Therefore, the North Paci c Ocean is an ideal natural laboratory for eolian dust All Rights Reserved. research. Numerous studies of eolian dust have focused on short-term dust deposition in the North Pacific

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Ocean. For instance, quartz abundance is used as an eolian proxy and is coherent with the δ18O record, where more quartz entered the ocean during glacial periods at core RC14-105 (Pisias & Leinen, 1984). Grain size fluc- tuations for eolian components in core KK75-02 were found to be controlled by orbital forcing over the past 700,000 years (Janecek & Rea, 1985). A Quaternary eolian record has also been constructed from core V21-146 on Shatsky Rise, where the maxima and minima of eolian fraction mass accumulation rates correspond exactly to loess and paleosol layers, respectively, in the Xifeng profile, Chinese Loess Plateau (Hovan et al., 1989). The history of eolian sedimentation over the past 70 million years has also been reconstructed from long cores (e.g., LL44-GPC3 and DSDP Site 576) by Janecek and Rea (1983) and Janecek (1985). A series of deep-sea sediment cores along a transect in the North Pacific Ocean was obtained during Ocean Drilling Program (ODP) Leg 145, which provided the opportunity to investigate the history of Cenozoic terri- genous deposition in the North Pacific Ocean. For example, about 70 m of pelagic post-Cretaceous sediment was cored at ODP Sites 885/886. This site is located downwind of Asia and is sufficiently far from the conti- nent to preclude riverine inputs and ice-rafted debris influences (Snoeckx et al., 1995). Based on isotope analyses, sediments from Site 885/886 have a provenance from the northern Tibetan Plateau and Gobi Desert (Pettke et al., 2000). Snoeckx et al. (1995) recovered a history of eolian deposition over the past 80 million years at Hole 885A. Sun and Liu (2007) reported a new higher-resolution eolian record from this core that spanned the late Pliocene to early Pleistocene and compared it with eolian deposition on the Chinese Loess Plateau. These studies have made significant contributions to understand eolian evolutionary history at different timescales. For these previous studies, the eolian dust signal preserved in pelagic sediments was constructed mainly by

chemical separation procedures, which made use of acetic acid, NaOH (stronger) or Na2CO3 (milder), and sodium dithionite to remove calcium carbonate, opal, and oxides (hydroxides), respectively. However, such chemical procedures are relatively time-consuming and are less suitable for high-resolution studies. Environmental magnetism has also been used to trace eolian signals on the basis of the magnetic properties (type, concentration, and grain size) of wind-borne minerals (Evans & Heller, 2003; Liu et al., 2012; Thompson & Oldfield, 1986). For example, Yamazaki and Ioka (1997) reported that biogenic magnetite produced in situ and detrital magnetite and hematite coexisted in red clay sediments in the North Pacific Ocean. Arnold et al. (1995) analyzed long-term magnetic variations at low resolution on sediments from ODP Sites 885 and 886, and Bailey et al. (2011) also reported magnetic parameters for Hole 885A. Properties of magnetic minerals contained in eolian sediment fractions are primarily controlled by source mineralogy, transportation path- ways, and water regime (Schwertmann & Taylor, 1987). Magnetic parameters of use in dust reconstructions include magnetic susceptibility (deMenocal et al., 1991), hard isothermal remanent magnetization (HIRM; Bloemendal et al., 1988; Doh et al., 1988; Robinson, 1986), and isothermal remanent magnetization (IRM) after

alternating field (AF) demagnetization at 120 mT (IRM900mT@AF120mT; Larrasoaña et al., 2003). However, multiple provenance and post-depositional processes, such as low-temperature oxidation (Johnson & Merrill, 1972; Van Velzen & Dekkers, 1999), reductive dissolution (Roberts, 2015), formation of biogenic magnetite (Hesse, 1994; Roberts, Florindo, et al., 2011), and volcanogenic inputs can cause ambiguities when interpreting the paleoenvironmental significance of magnetic proxies. Diffuse reflectance spectroscopy (DRS) is an alternative way to quantify both hematite and goethite contents (Deaton & Balsam, 1991) and has been used widely in soils (Balsam et al., 2004; Ji et al., 2004; Torrent et al., 2007), aerosols (Shen et al., 2006), and sediments (Balsam & Deaton, 1996; Liu et al., 2011; Oldfield et al., 2014; Zhang et al., 2007). DRS technology has been used, for example, to identify a combination of hematite and goethite in eolian sediment fractions derived from northwest Africa (Balsam et al., 1995). Jiang et al. (2013) and Hu et al. (2016) demonstrated that the combination of DRS and magnetic methods should provide a more reliable approach for determining the concentration of antiferromagnetic minerals in sediments. Previous magnetic studies of Hole 885A have focused on magnetic parameter variations, with less attention paid to the composition and morphology of magnetic minerals at the microscopic scale. Most importantly, there is no demonstrably reliable eolian dust proxy for this hole, which can preclude the influence of other sediment components. Despite demonstration of intensification of Northern Hemisphere glaciations and increased eolian dust flux to the North Pacific Ocean since the late Pliocene revealed by geochemistry and magnetic susceptibility (Haug et al., 2005; Maslin et al., 1998), little attention has been paid to changes in the type, concentration, and grain size (or domain state) of eolian magnetic minerals over this special

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climate transition. Compared with studies of ferrimagnetic minerals (magnetite and maghemite), hematite and goethite have not been well studied. ODP Hole 885A has been well investigated with various methods, which provide a solid paleoclimate background for comparison among different proxies. Based on integrated approaches, we aim to (1) extract wind-borne magnetic minerals from pelagic sediments at ODP Hole 885A; (2) assess changes in type, concentration, and grain size (or domain state) of ferrimagnetic minerals and analyze the corresponding relationship with the history of eolian dust deposition; and (3) identify and quantify hematite/goethite content variations to determine a more reliable eolian dust proxy for the North Pacific Ocean.

2. Materials and Methods In this study, 479 discrete samples were taken from sections 2H-1W to 2H-6W of ODP Hole 885A (44°410N, 168°160W; water depth = 5,708.5 m; Figure 1) between depths of 5.44 and 16.66 mcd (meters composite depth). These samples are composed predominantly of reddish brown and brown clay (with clay concentra- tions of 65–95%; Rea et al., 1993). Sediment age was obtained by linear interpolation on the basis of a magnetic reversal stratigraphy for Hole 885A (Dickens et al., 1995; Weeks et al., 1995). Age control points include the following: Brunhes/Matuyama boundary = 4.86 mcd, top Jaramillo = 5.70 mcd, top Olduvai = 9.57 mcd, onset Olduvai = 10.56 mcd, Matuyama/Gauss = 14.80 mcd, Gauss/Gilbert = 23.70 mcd. Hysteresis loops were measured to maximum applied fields of 1.5 T for samples using a vibrating sample

magnetometer (VSM, Princeton MicroMag 3900). Saturation magnetization (Ms), saturation remanent magne- tization (Mrs or SIRM1.5T), and coercivity (Bc) were obtained after high-field slope correction. The remanence coercivity (Bcr) was obtained by direct current demagnetization of SIRM1.5T back to 1.5 T. The high-field susceptibility (χp) was obtained from the slope of hysteresis loops fitted to data from 0.7 T to 1.4 T. The corrected susceptibility χferri = χlf χp reflects the contributions of ferrimagnetic minerals by removing the influence of paramagnetic minerals (χlf corresponds to the low-frequency susceptibility; Liu et al., 2012). IRM acquisition curves with 100 measurement points and nonlinear field steps on a log scale up to 1.5 T were measured using the VSM. In order to attenuate measurement noise, we used three-point running means to process the measured data. Then IRM acquisition curves were decomposed into magnetic components using the method of Kruiver et al. (2001). First-order reversal curves (FORCs) with 100 FORCs were also measured using the VSM for representative samples. FORC diagrams were processed using the FORCinel software of Harrison and Feinberg (2008). The S ratio and L ratio were obtained from the following procedure: a satura-

tion isothermal remanent magnetization was imparted in a 1.5 T applied field, termed SIRM1.5T, then samples were demagnetized with backfields of 100 and 300 mT, and the corresponding remanence were mea-

sured, termed IRM100mT and IRM300mT. HIRM300T is defined as 0.5 × (SIRM1.5T + IRM300mT), and the S ratio is defined as IRM300mT/SIRM1.5T (King & Channell, 1991). Analogously, HIRM100mT is defined as 0.5 × (SIRM1.5T + IRM100mT). HIRM100mT and HIRM300mT are equivalent to the SIRM after AF demagnetization at 100 (IRM@100mT) and 300 mT (IRM@300mT), respectively (Liu et al., 2007), and the L ratio is defined as HIRM300mT/HIRM100mT (Liu et al., 2007). Low (470 Hz) and high (4700 Hz) frequency magnetic susceptibility and anhysteretic remanent magnetization susceptibility (χARM) data were obtained from Bailey et al. (2011). In order to determine the magnetic mineralogy, low-temperature experiments were conducted using a Quantum Design Magnetic Property Measurement System for representative samples. First, samples were

cooled from 300 to 20 K in zero field, then a SIRM2.5T was imparted at 20 K in a 2.5 T field. The SIRM2.5T was then demagnetized thermally by warming from 20 to 300 K in zero field at a temperature sweep rate of 5 K/min. Temperature dependent susceptibility (χ-T) curves were measured for representative samples using a Kappabridge MFK1-FA magnetic susceptibility meter in an argon environment. Magnetic minerals were extracted from bulk sediment for selected samples and were deposited onto carbon- coated copper grids for transmission electron microscopy (TEM) investigations. TEM experiments were performed with a JEOL 2100 transmission electron microscope operating at 200 kV. Energy-dispersive X- ray spectroscopy (EDXS) analysis was performed to analyze the chemical composition of targets at nanometer scales. DRS was measured for all samples using a Cary 5000 ultraviolet-visible-infrared spectrometer equipped

with BaSO4 as the white standard. The scan rate was 300 nm/min from 300 to 2,600 nm in 0.5 nm

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Figure 1. Location of ODP Hole 885A (red circles), Asian inland dust sources (gray white shadow area with number 1 to 6), eolian deposition area in North Pacific Ocean (light-medium-dark yellow area), and trajectories of westerlies (green arrows; Merrill et al., 1989). Asian dust source numbers 1 to 6 indicate Taklimakan, Gobi, Badain Juran, Tengger, Mu Us, Mongolia Gobi desert, or sandy desert, respectively (Sun et al., 2013). “X” with dashed black line indicates southern limit ice-rafted debris (Bigg et al., 2008). The dashed and purple lines indicate the fluxes of mineral aerosol to the oceans (in mg cm 2 ka 1; Duce et al., 1991). The black number in contours within yellow area indicates spatial distribution of <2 μm fraction weight percent of surface samples in the North Pacific Ocean (Arnold, 1996); contour interval is 5%.

steps. DRS data were first transformed into the Kubelka-Munk (K-M) remission function (F(R)=(1 R)2/2R, R is reflectance), and then second derivative curves of F(R) were calculated following Scheinost et al. (1998) and Torrent et al. (2007). Characteristic bands are located at approximately 535 and 425 nm for pure hematite and goethite, respectively, in second derivative curves (Scheinost et al., 1998; Torrent et al.,

2007). The corresponding band intensity for hematite and goethite are defined as IHm and IGt, which are proportional to the concentration of hematite and goethite (Scheinost et al., 1998). We have developed a new proxyRelHm+Gt, which represents the relative concentration of antiferromagnetic I I P Hmi P Gti minerals: RelHmþGt ¼ n n þ n (IHm and IGt indicate the intensity of hematite and goethite, i I I i¼1 Hmi i¼1 Gti and n denotes the number of samples).

3. Results 3.1. Variation of Magnetic Parameters With Depth Environmental magnetic parameters for ODP Hole 885A are shown in Figure 2 from 2.79 to 0.93 Ma. On the basis of trends in these parameters, the section can be divided into three units: Unit 1: 2.58–2.79 Ma, Unit 2: 2.03–2.58 Ma, and Unit 3: 0.93–2.03 Ma. Results are discussed below for each unit, respectively.

Unit 1 (2.58–2.79 Ma): Variations of χferri, χp, χARM, Ms, and SIRM1.5T suggest that the concentration of ferrimag- netic (magnetite/maghemite) and paramagnetic (silicate) minerals decreases first and then gradually

increases upward to the transitional zone (at 2.58 Ma) between Units 1 and 2. Generally, χARM/SIRM variations reflect changes in magnetic grain size (Peters & Dekkers, 2003; Thompson & Oldfield, 1986); however, ARM

may be influenced significantly by interactions among magnetic grains (Sugiura, 1979), so that χARM/ SIRM1.5 may not indicate accurately grain size changes at Hole 885A. Samples from Unit 1 are characterized by higher χARM/SIRM1.5T, which may reflect weaker magnetic interactions where particles acquire a higher ARM. Bcr is sensitive to magnetic grain-size and mineralogy changes. The increase in concentration- dependent magnetic parameters at ~2.58 Ma is associated with a synchronous change in χARM/SIRM1.5T and Bcr values, which indicates that magnetic interactions and high-coercivity magnetic minerals increased at this point. χfd is an indicator of fine magnetic particles near the superparamagnetic (SP)/single domain (SD) threshold (Dearing et al., 1996; Liu et al., 2012). Thus, the lower χfd value in this unit suggests an insignif- icant contribution of SP particles. The HIRM and S ratio can be used to quantify the absolute and relative concentration of antiferromagnetic minerals (hematite and goethite), respectively, only when the L ratio is relatively constant (Liu et al., 2007). These two proxies cannot be used to interpret changes in hematite and goethite contents because of drastic L ratio fluctuations in this unit. In the transition zone between

Units 1 and 2, the L ratio is relatively stable; hence, the change in HIRM300mT value likely indicates

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increased hematite and goethite inputs at 2.58 Ma, which is consistent with previous reports from the low-latitude North Pacific Ocean (Yamazaki & Ioka, 1997).

Unit 2 (2.03–2.58 Ma): Overall, the values of χferri, Ms, SIRM1.5T, and HIRM300mT are relatively higher between 2.03 and 2.58 Ma than in Unit 1. χp is relatively constant, which indicates a stable input of silicates, while χARM and the S ratio decrease, χARM/SIRM1.5T is low, and Bcr increases remarkably compared to Unit 1. It can be inferred that the concentration of particles with stronger magnetic interactions and higher coercivity increases in this unit. Unit 3 (0.93–2.03 Ma): In this unit, four volcanic ash layers (Rea et al., 1993)

have high concentration-dependent parameters and Bcr. Apart from these layers, χferri, χp, Ms, SIRM1.5T, HIRM300mT, and Bcr are all higher than in Unit 2. In general, strongly similar variations among concentration-dependent

magnetic parameters (χferri, Ms, SIRM1.5T, and HIRM300mT) suggest that magnetic enhancement across the transition at 2.58 Ma is dominantly con- trolled by increased ferrimagnetic (magnetite and maghemite) and anti- ferromagnetic mineral (hematite and goethite) contents. 3.2. Hysteresis Data

Plots of Mrs/Ms and Bcr/Bc can provide information about magnetic prop- erty variations in a sample set (Day et al., 1977; Dunlop, 2002a, 2002b). Data from surface samples from Asian inland dust sources (Taklimakan Desert and Qaidam Basin; Liu et al., 2015) extend the trend from the distri-

bution of data points for ODP Hole 885A to lower Mrs/Ms and higher Bcr/Bc values that are indicative of coarser multidomain particles (Figure 3a). The data distribution from finer-grained material at Hole 885A to coarser in Asian inland dust sources (Figure 3a) is interpreted as a response to sorting by wind from source to sink. The overall data trend is from the upper left to lower right, where data for volcanic ash layers diverge from the trend due to abundant SP particles in the ash layers (Figure 3b). 3.3. Bivariate Diagrams for Magnetic Parameters The studied samples are predominantly reddish brown and brown clay (Rea et al., 1993), so the influence on magnetic properties of post- Figure 2. Magnetic parameters versus age of Hole 885A from 2.79 to depositional magnetic mineral oxidation should be considered, especially 0.93 Ma. (a) Ferrimagnetic susceptibility, χferri = χlf χp (χlf, low-frequency transformations from SD magnetite to SP maghemite through oxidation in χ χ susceptibility; p, paramagnetic susceptibility); (b) fd, frequency-dependent pelagic clays (Johnson & Merrill, 1972; Kent & Lowrie, 1974). The absence of susceptibility (χfd = χlf χhf; χlf, low-frequency susceptibility; χhf, correlation between χferri and χfd for non-ash layers may reflect that varia- high-frequency susceptibility), (c) χp, paramagnetic susceptibility; (d) χARM, tions in the amount of SP particles are independent of the total amount of anhysteretic remanent magnetization susceptibility; (e) Ms, saturation magnetization; (f) SIRM1.5T, saturation isothermal remanent magnetization; magnetic minerals (Figure 4a). In contrast to samples without volcanic (g) HIRM300mT, hard remanent magnetization; (h) L ratio; (i) S ratio; (j) Bcr, influences, there is a positive relationship between these two parameters χ remanence coercivity; and (k) ARM/SIRM1.5T. The shallow, medium, and in volcanic ash layers, which indicates that volcanism has generated abun- deep gray shading indicate Unit 1, Unit 2, and Unit 3, respectively. The four fi χ vertical dotted bands depict volcanic ash layers (labeled as Ash 1, Ash 2, dant ultra ne SP particles (Figure 4a). ferri is directly proportional to Ms Ash 3, and Ash 4; Rea et al., 1993). Note that the value of χARM, Ms, SIRM1.5T, (Figure 4b), which demonstrates that the bulk magnetic properties are and HIRM300mT were clipped at volcanic ash layers in order to clarify the controlled mainly by the ferrimagnetic mineral concentration. Indicators variation tendency. The original value of related magnetic parameters is of ferrimagnetic minerals (χferri) and silicates (χp) fall into two groups exhibited in Figure S1 in the supporting information. (Figure 4c). These two parameters are correlated positively especially in Group 1, which suggests that the concentration of ferrimagnetic minerals

and silicates could have a similar origin. HIRM300mT and L ratio have a linear trend for samples in Unit 1 (Figure 4d). Therefore, HIRM300mT variations are affected largely by hematite/goethite coercivity in this unit. Ms for magnetite/maghemite is about two orders of magnitude higher than for hematite/goethite and is

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Figure 3. (a) Day plot for samples from Hole 885A, Qaidam Basin, and Taklimakan desert (the latter two are from Liu et al., 2015) and (b) magnified Day plot, dashed green circle separates the samples in volcanic ash layers from other samples.

independent of grain size; hence, Ms provides a measure of the absolute concentration of ferrimagnetic minerals. HIRM300mT is correlated positively with Ms (Figure 4e), which suggests that hematite/goethite and ferrimagnetic mineral have a similar origin.

3.4. Temperature-Dependent Measurements Temperature-dependent magnetic susceptibility (χ-T) measurements can provide information about miner-

alogy and transformation of magnetic minerals by recognizing the Curie temperature (Tc) and other special features (Dunlop & Özdemir, 1997). χ-T curves for selected samples are shown in Figure 5 (Figures 5a-1, 5a-5,

Figure 4. Bivariate diagrams of magnetic parameters. Correlation diagrams between (a) χferri and χfd, (b) χferri and Ms, (c) χferri and χp, (d) HIRM300mT and L ratio, and (e) HIRM300mT and Ms.

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Figure 5. Temperature-dependent measurement curves, FORC diagrams, and decomposition of IRM acquisition curves. (a-1, a-5, b-1, b-5, c-1, c-5, d-1) Curves of tem- perature-dependent magnetic susceptibility; (a-2, a-6, b-2, b-6, c-2, c-6, d-2) Low-temperature warming of SIRM curves; (a-3, a-7, b-3, b-7, c-3, c-7, d-3) FORC diagrams, with the smoothing factor of 3; (a-4, a-8, b-4, b-8, c-4, c-8, d-4) decomposition of IRM acquisition curves; the pink, green, deep blue, light blue, and red curves in decomposition diagrams denote Component 1, Component 2, Component 3, Component 4, and sum of components, respectively.

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5b-1, 5b-5, 5c-1, 5c-5, and 5d-1). All heating curves (red) are characterized by a sharp drop at about 585°C,

which is the Tc of magnetite (Evans & Heller, 2003). However, the cooling curves lie above the heating curves due to formation of new magnetite from other minerals (Fe-bearing silicates and clay minerals) during heating (Dunlop & Özdemir, 1997; Hunt et al., 1995). Low-temperature measurements have been widely used to identify mineralogy and grain size (or domain state) of magnetic minerals (Moskowitz et al., 1998; Özdemir et al., 1993). For example, the Verwey (Verwey, 1939) and Morin transitions (Morin, 1950) are key features for identifying magnetite and hematite,

respectively. For stoichiometric magnetite, the Verwey transition temperature (Tv) is near 120 K, and the amplitude of magnetization decrease at Tv is gradually suppressed when domain state changes from the multidomain to the SD state (Dunlop & Özdemir, 1997) because magnetite is usually partially oxidized to

maghemite, so that Tv can be partially or completely suppressed in fine particles (Özdemir et al., 1993). All samples have almost identical curves with no obvious Verwey transition (Figures 5a-2, 5a-6, 5b-2, 5b-6, 5c-2, 5c-6, and 5d-2). This suggests that magnetite particles in the studied samples have undergone significant oxidation, as is the case with other pelagic sediments from the Pacific Ocean (Smirnov & Tarduno, 2000; Yamazaki, 2009; Yamazaki & Ioka, 1997).

3.5. FORC Diagrams and IRM Acquisition Curves FORC diagrams can be used to identify domain state and to discriminate among different components in a magnetic mineral assemblage (Pike et al., 1999; Roberts et al., 2000). In Unit 1, FORC diagrams are character-

ized by elongated contours along the horizontal axis with small vertical spread (Bu < 5 mT; Figures 5a-3 and 5a-7), which indicate typical SD behavior with weak magnetostatic interactions (Roberts et al., 2000). FORC diagrams for samples from Units 2 and 3 (Figures 5b-3, 5b-7, 5c-3, and 5c-7) have diverging outer contours that are indicative of vortex state particles in addition to the non-interacting SD particles indicated by the inner contours (Roberts et al., 2000, 2017). A FORC diagram for a volcanic ash layer contains a secondary peak near the origin of the diagram, with contours nearly parallel to the vertical axis (Figure 5d-3), which is indica- tive of SP particles (Pike et al., 2001; Roberts et al., 2000), along with a main peak due to SD and vortex state

particles. Along the Bc axis, the horizontal contours extend to high fields (i.e., 80–100 mT; Figure 5d-3), which suggests the presence of abundant high-coercivity minerals in volcanic ash layers. IRM acquisition curves can be decomposed into coercivity fractions (Heslop et al., 2002; Kruiver et al., 2001) to provide key information for discriminating components with different magnetic properties. In general, four

components are recognized: Component 1 with B1/2 of 10–18 mT and dispersion parameter (DP) of 0.24–0.29; Component 2 with B1/2 of 35–42 mT and DP of 0.20–0.22; Component 3 with B1/2 of 85–141 mT and DP of 0.19–0.28; and Component 4 with B1/2 of 260–380 mT and DP of 0.17–0.21. The details are summar- ized in Table 1. Component 1 may represent ultrafine SP magnetite (maghemite) just above the SP/SD thresh- old size (~20–25 nm), which can have SD behavior because the measurement time is less than the thermal relaxation time (Heslop et al., 2004). Component 2 has a relatively constant mean coercivity and narrow DP and is the major magnetic carrier in Unit 1 (Figures 5a-4 and 5a-8) and is present in all samples. Combined

with the corresponding FORC diagrams (Figures 5a-3 and 5a-7) and higher χARM/SIRM, this component probably represents biogenic SD magnetite, which is present in all samples and has similar magnetic proper- ties as the biogenic magnetite that occurs widely in other Pacific Ocean sediments (Yamazaki, 2009; Yamazaki & Shimono, 2013). Component 3 is characterized by higher coercivity and occurs in significantly higher concentrations in samples from Units 2 and 3 after 2.58 Ma (Figures 5b-4, 5b-8, 5c-4, and 5c-8). This compo- nent is interpreted as eolian maghemite in pelagic North Pacific Ocean sediments and is characterized by high coercivity (~100 mT) and stronger interactions (Yamazaki, 2008, 2009). However, Hu et al. (2013) found a fine pedogenic hematite (~130 mT) in loess/paleosol sequences of the Chinese Loess Plateau. In addition, partially oxidized magnetite can also have elevated coercivity due to the mismatching of unit cells between the maghemite rim and the magnetite core (Cui et al., 1994; Özdemir & O’Reilly, 1982). Component 4 contri- butes <10% to the bulk SIRM and is attributed to antiferromagnetic hematite/goethite.

3.6. TEM Observations TEM observations are used widely to identify magnetic minerals preserved in sediments (Chang et al., 2016; Fassbinder et al., 1990; Maher, 1988). TEM observations from the studied sediments indicate abundant

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Table 1 The Results of Decomposition of IRM Acquisition Curves for Selected Samples at Hole 885A Component 1 Component 2 Component 3 Component 4

IRM IRM IRM Sample Age IRM B1/2 IRM IRM B1/2 Rel. IRM B1/2 Rel. IRM B1/2 Rel. ID (Ma) Abs. (mT) DP Rel. (%) Abs. (mT) DP (%) Abs. (mT) DP (%) Abs. (mT) DP (%)

885A-2H1W-77-78 1.08 3.90E03 15.49 0.24 14.89 1.10E02 39.81 0.21 41.98 9.80E03 100.00 0.25 37.40 1.50E03 323.59 0.20 5.73 885A-2H1W-101.5-102.5 1.14 3.70E03 17.78 0.28 18.32 9.00E03 40.74 0.22 44.55 6.00E03 100.00 0.22 29.70 1.50E03 269.15 0.21 7.43 885A-2H2W-12-13 1.27 2.60E03 15.85 0.27 18.51 6.45E03 38.90 0.21 45.91 4.45E03 104.71 0.25 31.67 5.50E04 354.81 0.20 3.91 885A-2H2W-71-72 1.40 6.30E03 15.49 0.25 14.35 1.59E02 38.02 0.21 36.22 1.90E02 93.33 0.25 43.28 2.70E03 331.13 0.20 6.15 885A-2H2W-114-115 1.48 4.00E03 17.78 0.28 21.05 8.50E03 40.74 0.22 44.74 5.00E03 100.00 0.22 26.32 1.50E03 269.15 0.21 7.89 885A-2H3W-21-22 1.57 4.90E03 15.85 0.24 14.80 1.35E02 39.81 0.21 40.79 1.25E02 102.33 0.25 37.76 2.20E03 316.23 0.20 6.65 885A-2H3W-65-66 1.66 2.70E03 16.98 0.27 16.29 6.50E03 39.81 0.21 39.23 6.60E03 100.00 0.27 39.83 7.70E04 380.19 0.20 4.65 885A-2H3W-104-105 1.74 6.80E03 17.78 0.28 18.94 1.42E02 41.69 0.22 39.55 1.10E02 100.00 0.22 30.64 3.90E03 269.15 0.21 10.86 885A-2H4W-45-46 1.96 6.00E03 15.14 0.25 13.25 1.90E02 38.90 0.21 41.94 1.70E02 97.72 0.24 37.53 3.30E03 295.12 0.20 7.28 885A-2H4W-62-63 1.99 1.60E02 17.78 0.28 17.80 3.63E02 40.27 0.22 40.38 2.80E02 100.00 0.22 31.15 9.60E03 263.03 0.21 10.68 885A-2H4W-101-102 2.12 2.70E03 17.78 0.28 22.98 5.50E03 38.46 0.22 46.81 2.60E03 100.00 0.21 22.13 9.50E04 275.42 0.20 8.09 885A-2H5W-15-16 2.32 2.30E03 15.85 0.26 21.18 5.50E03 38.02 0.20 50.64 2.80E03 109.65 0.28 25.78 2.60E04 380.19 0.20 2.39 885A-2H5W-42-43 2.40 2.80E03 15.85 0.29 21.37 6.80E03 38.02 0.21 51.91 3.00E03 112.20 0.25 22.90 5.00E04 354.81 0.20 3.82 885A-2H5W-76-77 2.50 5.00E03 17.78 0.28 17.18 1.35E02 41.69 0.22 46.39 7.70E03 100.00 0.21 26.46 2.90E03 260.02 0.20 9.97 885A-2H5W-96-97 2.55 5.90E03 16.60 0.25 18.50 1.30E02 38.90 0.20 40.75 1.16E02 85.11 0.25 36.36 1.40E03 346.74 0.20 4.39 885A-2H5W-103-104 2.57 5.00E03 17.78 0.28 17.41 1.56E02 41.21 0.22 54.39 6.50E03 100.00 0.21 22.63 1.60E03 263.03 0.20 5.57 885A-2H6W-3-4 2.65 5.50E03 12.59 0.29 18.39 1.95E02 38.90 0.22 65.22 3.20E03 114.82 0.20 10.70 1.70E03 295.12 0.20 5.69 885A-2H6W-65-66 2.74 8.30E04 10.00 0.28 12.67 4.70E03 35.48 0.21 71.76 7.00E04 112.20 0.20 10.69 3.20E04 316.23 0.20 4.89 885A-2H6W-77-78 2.76 1.20E03 13.65 0.26 15.79 5.50E03 38.02 0.22 72.37 7.20E04 141.25 0.19 9.47 1.80E04 363.08 0.17 2.37 2 Note. IRM Abs. denotes the absolute value of isothermal remanent magnetization (Am /kg), B1/2 is the field at which half of the SIRM is reached, DP is dispersion parameter, and IRM Rel. denotes the relative percentage content of isothermal remanent magnetization for each component. 10.1002/2017JB014951 3366 Journal of Geophysical Research: Solid Earth 10.1002/2017JB014951

magnetic minerals with a wide range of grain size and diverse morphologies and can be classified into four types as described below. The first magnetic mineral type is interpreted to be biogenic magnetite (Figures 6a–6d), which has octahedral morphologies and grain size of several tens of nanometers. It is well known that magnetosomes occur abun- dantly in Pacific Ocean pelagic red clays (Yamazaki & Ioka, 1997; Yamazaki & Shimono, 2013). FORC diagrams and IRM acquisition analysis indicate the presence of biogenic magnetite at Hole 885A, which is confirmed by TEM observations. The second magnetic mineral type identified with TEM is nano-sized magnetic mineral inclusions embedded within host silicates, which are present in most selected samples (Figures 6e–6h). The main inclusion micro- structure is characterized by dendritic clusters (Figure 6g) and dendrites (Figures 6e, 6f, and 6h) with mean grain size <100 nm, which is approximately within the size range for SD magnetite. Selected area electron diffraction analysis indicates that these magnetic inclusions have high crystallinity and clear diffraction pat- terns (Figure 6o). EDXS analysis reveals that the inclusions contain abundant Fe and O, with a small quantity of Ti, Cr, Mg, and Co (Figures 6q-2 and 6r-2). The host minerals are rich in O and Si and also have small Al and K contents (Figures 6q-1 and 6r-1). EDXS results for inclusions also indicate Si (e.g., Figure 6q-2) because the inclusions are embedded in host silicate minerals. All of the above morphological, crystallographic, and che- mical compositional characteristics demonstrate that the inclusions are nano-sized (titano) magnetite and that the hosts are likely silicates (e.g., feldspar or pyroxene). The nano-sized magnetic inclusions are adjacent to each other, which may cause significant magnetic interactions that influence ARM values. Magnetic inclu-

sions occur within silicate hosts, so that high-field magnetic susceptibility (χp) indicates magnetic inclusions concentration changes. The third type of magnetic mineral particle identified with TEM occurs mainly in Unit 2 sediments and has submicron scale (>200 nm), subangular morphology (Figures 6k and 6l), and good crystallinity (Figure 6p). This component may be coarser detrital magnetite, which brings vortex state signatures into corresponding FORC diagrams (Figures 5b-3, 5b-7, 5c-3, and 5c-7). In addition to the above three types of magnetic particles, ultra-fine SP maghemite aggregates (Figure 6j) are recognized on the basis of polycrystalline morphological features (Figure 6i). The aggregate maintains the profile of subangular primary coarser vortex state magnetite (Figure 6k), which suggests that SP maghemite might have formed by oxidation of magnetite (Johnson & Merrill, 1972; Kent & Lowrie, 1974).

3.7. Summary of Ferrimagnetic Minerals at Hole 885A Based on the above magnetic and TEM results, the magnetic minerals preserved in the three units of Hole 885A have distinct features as follows: Unit 1 (2.58–2.79 Ma): Based on FORC diagrams, IRM acquisition analysis, and TEM images, biogenic magnetite with minimal magnetostatic interactions is the dominant magnetic carrier in Unit 1 and represents a magnetic background component in all samples. Unit 2 (2.03–2.58 Ma): Coexisting with the biogenic magnetite background, nano-sized (titano) magnetite inclusions

in silicate hosts have significantly increased since 2.58 Ma, as inferred from the increased χp between Units 1 and 2. The relative contribution of Component 3 from IRM analysis also increased since 2.58 Ma (Figures 5b-4, 5b-8, 5c-4, and 5c-8). The fine (titano) magnetite inclusions fall within the SD particle size range and are partially substituted by Ti, Cr, Mg, and Co (Figures 6q-2 and 6r-2). Therefore, this component may have higher coercivity than stoichiometric magnetite. Such inclusions are also adjacent to each other, which may produce

strong magnetic interactions (lower χARM/SIRM). This evidence suggests that Component 3 may represent magnetic inclusions. However, partially oxidized magnetite and fine pedogenic hematite may have coercivity distributions similar to Component 3. Our experimental results may not be enough to identify this compo- nent. Regardless, Component 3 is likely to be of eolian origin. In addition to the above fine particles, coarser detrital vortex state magnetite also increased in abundance at around 2.58 Ma as indicated in FORC diagrams and TEM images. Part of the vortex state magnetite has been transformed into ultrafine SP maghemite, which occurs as polycrystalline aggregates with profiles similar to subangular magnetite, which suggests that it is a product of post-depositional oxidation of magnetite. Unit 3 (0.93–2.03 Ma): Volcanic activity generated abun- dant magnetic minerals and silicates in this unit. Volcanogenic particles are characterized by stronger magnetism and higher coercivity. Regardless of these ash layers, the content of (titano) magnetic inclusions further increased in Unit 3 compared with Unit 2.

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Figure 6. Transmission electron microscopy (TEM) images, selected area electron diffraction patterns, and energy-dispersive X-ray spectroscopy (EDXS) analysis. (a-h, j-l) TEM images for magnetic extracts (a, e, f) from samples ODP-885A-2H6W-65-66, (b, d, g, h) from sample ODP-885A-2H6W-23-24, (c, j, k, l) from sample ODP-885A- 2H4W-147-148, (m–p) SAED, (i) polycrystalline diffraction ring for (j), (q, r) EDXS spectra for magnetic inclusions (q-2 and r-2) and their host minerals (q-1 and r-1). The purple cross sign in Figures 6a, 6d, 6h, and 6l indicate the selected area for SAED analysis. Blue and red crosses in panels 6e and 6g indicate the EDXS analysis spots for hosts and inclusions, respectively. The Cu peaks in panels 6q and 6r originate from the TEM grid. Values of Miller indices (hkl) are indicated in panels 6m–6p, which are consistent with those of titanomagnetite or magnetite. The details of the magnetic extraction procedures are exhibited in Text S1 in the supporting information.

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Figure 7. Diffuse reflectance spectroscopy (DRS) results. (a, b) Intensity of goethite and hematite obtained from second derivative of K-M remission function from DRS, (c) RelHm+Gt, (d) percentage content of eolian fraction separated by a chemical method from Sun and Liu (2007), (e) HIRM300mT for studied samples at Hole 885A. The shallow, medium, and deep gray shadings indicate Unit 1, Unit 2 and Unit 3, respectively. The four horizontal dotted bands depict the volcanic ash layers (labeled as Ash 1, Ash 2, Ash 3, and Ash 4; Rea et al., 1993). Note that the value of HIRM300mT was clipped at volcanic ash layers in order to clarify the variation tendency. The original value of HIRM300mT is exhibited in Figure S2 in the supporting information.

3.8. DRS Results

Temporal variations of goethite (IGt) and hematite (IHm) intensity obtained from second derivative curves of the DRS K-M remission function, RelHm+Gt, the percentage content of the eolian fraction separated by the chemical method of Sun and Liu (2007), and HIRM300mT are shown in Figure 7. There is a strong similarity in variations of RelHm+Gt (Figure 7c) and the percentage content of eolian dust (Figure 7d), especially in the transition between Units 1 and 2, where they increase synchronously at 2.58 Ma and rise gradually from 1.50 to 0.93 Ma. However, the intensity curves for goethite (Figure 7a) and hematite (Figure 7b) are relatively noisy and do not mirror variations for the eolian component. Eolian sediment fractions often contain a combination of hematite and goethite (Balsam et al., 1995). Liu et al. (2015) inferred that hematite and goethite in surface dust samples in Taklimakan Desert, Qaidam Basin, and Mongolia-Gobi Desert could have

a similar provenance (Figure 8a). IHm and IGt for Hole 885A have a similar distribution pattern (Figure 8b) compared to dust from Asian source regions (Figure 8a), which suggests that hematite and goethite at

Hole 885A are derived from Asian dust sources. HIRM300mT variations (Figure 7e) are consistent with RelHm+Gt (Figure 7c) and percent eolian content (Figure 7d) in Units 1 and 2. However, HIRM300mT values in Unit 3 are influenced significantly by volcanic activity, especially for samples between Ash 2 and Ash 4.

4. Discussion 4.1. Ferrimagnetic Minerals and Eolian Dust Inputs Widespread arid to semiarid regions in northwestern China provide abundant eolian sediment to the North Pacific Ocean by prevailing westerlies (Merrill et al., 1989). Increased ice-rafted debris in deep-sea sediments from the Pacific and Atlantic Oceans reflect intensification of Northern Hemisphere glaciations between 2.74 and 2.54 Ma (Maslin et al., 1998). During this period, the eolian sediment flux to the North Pacific Ocean increased significantly, which indicates intensified aridification in dust source regions and enhanced

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Figure 8. Correlation diagrams between intensity of goethite and hematite for samples from (a) Asian inland dust source (data from Liu et al., 2015) and (b) ODP Hole 885A.

atmosphere circulation (Janecek & Rea, 1983; Leinen, 1985). The studied pelagic red clay mainly consists of biogenic, terrigenous, and volcanogenic particles. Increased volcanogenic inputs from the Kamchatka and Aleutian Arcs (Prueher & Rea, 2001) played an important role in enhancing magnetic mineral concentrations in North Pacific Ocean sediments, which produce ambiguities when interpreting eolian signals using magnetic proxies alone. Nevertheless, by combining different magnetic properties, environmental magnetism still provides key information for discriminating sediment components with different origins and for evaluating the degree of volcanogenic particle influences. Hysteresis parameters (Figure 3b) can be an effective tool for discriminating between volcanogenic and terrigenous components and reveal that volcanic ash layers contain abundant high coercivity magnetic minerals with different magnetic behavior from eolian magnetite/maghemite derived from Asia. Environmental magnetism analyses indicate a weaker influence of volcanogenic particles in Units 1 and 2. The magnetic enhancement at 2.58 Ma is mainly controlled by eolian dust inputs from Asia with insignificant volcanogenic inputs. Yamazaki and Ioka (1997) reported that biogenic and detrital magnetite are the major ferrimagnetic compo- nents in North Pacific red clay. The former has lower coercivity and negligible magnetostatic interactions, while the latter has intermediate coercivity and stronger interactions (Yamazaki, 2009). Fine (titano) magne- tite inclusions embedded in silicates have been identified in most of the studied samples. The widespread occurrence of magnetite within silicates in deep-sea sediments has been reported by Chang et al. (2016). Isotopic analysis demonstrates that the silicate fraction at Site 885/886 over the past 11 Ma originated from Asian dust sources (Pettke et al., 2000), which indicates that the fine (titano) magnetite inclusions in silicates

were also derived from Asian inland dust sources. Paramagnetic susceptibility (χp) is often used as an indica- tor of silicates. In this study, χp indicates that magnetite inclusions increased significantly at 2.58 Ma, which suggests further enhanced aridity in Asian dust source regions under intensified Northern Hemisphere glaciations since the late Pliocene (Haug et al., 2005; Janecek & Rea, 1983; Shackleton et al., 1984; Snoeckx et al., 1995; Sun et al., 2006). Environmental magnetism and TEM observations indicate that the change in ferrimagnetic minerals could reflect increased eolian inputs to Hole 885A since the late Pliocene. However, magnetic parameters may

not extract precisely eolian signals because of biogenic and volcanogenic contributions. χp is an ambiguous indicator of eolian magnetite inclusions because volcanic activity also generates abundant silicates. Coarser eolian magnetite could be a more reliable eolian fraction indicator; however, it cannot be identified exclu- sively from magnetic parameters. Therefore, it is difficult to obtain a reliable eolian dust proxy from ferrimag- netic minerals alone. We, therefore, focus below on widespread and stable antiferromagnetic minerals.

4.2. RelHm+Gt From DRS as a Proxy for Eolian Dust Eolian dust typically contains a mixture of magnetically ordered iron oxides and oxyhydroxides, which are sensitive to the process of formation (Schwertmann & Taylor, 1987), transportation (Zhang & Iwasaka, 2004), and deposition (Bailey et al., 2011). However, ferrimagnetic minerals (magnetite/maghemite) are unstable in both oxidizing (Johnson & Merrill, 1972; Kent & Lowrie, 1974) and reducing (Roberts, 2015)

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sedimentary environments. In addition, magnetite/maghemite derived from various sources (e.g., dust, volcanic activity, ice-rafted debris, and magnetosomes) with different magnetic properties may cause ambi- guities in interpreting eolian signals using environmental magnetism. Hematite and goethite are volumetrically dominant magnetic minerals in rocks and sediments (Barron & Montealegre, 1986). Eolian dust from North Africa (Itambi et al., 2009; Larrasoaña et al., 2003, 2008; Lyons et al., 2010; Oldfield et al., 2014; Robinson, 1986), the Asian interior (Arnold et al., 1995; Bailey et al., 2011; Yamazaki, 2009; Yamazaki & Ioka, 1997), and Arabian Peninsula (Bloemendal et al., 1993; Hounslow & Maher, 1999) contains abundant antiferromagnetic minerals (hematite/goethite). Hematite is generally regarded as an end product of mineral weathering process (Barrón et al., 2003). Compared with ferrimagnetic minerals, hematite and goethite are relatively stable in various environments. Hence, antiferro- magnetic minerals are more useful as eolian indicators. HIRM is conventionally employed as a concentration-dependent proxy for hematite and goethite (Thompson & Oldfield, 1986). Doh et al. (1988) found that HIRM variations correlate remarkably well with eolian fluxes in North Pacific pelagic sediments. However, antiferromagnetic minerals may not reach saturation under laboratory magnetic fields, and the accuracy of HIRM may be influenced significantly by high coercivity ferri- magnetic minerals and variations in the coercivity distribution of hematite and goethite (Liu et al., 2007). For example, volcanic ash layers at Hole 885A have high HIRM values, which can be interpreted as due to abun-

dant hematite/goethite. However, the DRS record does not indicate distinct hematite (IHm) and goethite (IGt)

intensity changes within the ash layers. Inconsistencies between HIRM peaks and IHm and IGt minima for volcanic ash layers strongly indicate that, except for hematite and goethite, HIRM can be contaminated by magnetically hard minerals and that DRS parameters can efficiently attenuate the influence of volcanogenic components. Previous studies have demonstrated that high-Ti magnetite is the dominant iron oxide in tephra from Kamchatka (Braitseva et al., 1997; Ponomareva et al., 2015), the Aleutian Arc (Myers et al., 1985; Nye & Turner, 1990; Romick et al., 1990), and in Okhotsk Sea sediments (Wang et al., 2014). In addition, cation sub- stitution in the magnetite crystal lattice may increase coercivity significantly (Akimoto, 1962; Jiang et al., 2016; Liu et al., 2004). Abundant high-coercivity volcanogenic magnetite may disturb the original coercivity distri- bution, so that HIRM may not indicate the concentration of antiferromagnetic minerals in volcanic ash layers. DRS provides a quantitative means to determine the concentration of antiferromagnetic minerals (hematite and goethite) and has a lower detection limit than magnetic measurements, and it is a rapid and nondestruc- tive measurement (Deaton & Balsam, 1991). This method has been used extensively to extract hematite/goethite information in different natural settings (Balsam et al., 2004; Balsam & Deaton, 1996; Ji et al., 2004; Liu et al., 2011; Oldfield et al., 2014; Shen et al., 2006; Torrent et al., 2007; Zhang et al., 2007). The band location and amplitude intensity in DRS spectra are influenced by many factors, such as mathema- tical treatment methods, band overlap, and composition and concentration of minerals. We have adopted the second derivative curves of F(R)=(1 R)2/2R (R is reflectance) to obtain the concentration of hematite and goethite. Compared with first derivative curves, the band location in second derivative curves is not influ- enced by other constituents or by the concentration of minerals in mixtures, and the second derivative is more suitable for natural samples (Liu et al., 2011). Besides natural samples, previous studies have investi- gated synthetic hematite/goethite samples using DRS and traditional rock magnetic methods and found that the DRS band amplitude in second order derivative curves is a relatively reliable concentration proxy for

hematite and goethite (Hu et al., 2016; Jiang et al., 2013; Liu et al., 2011). In this study, we define RelHm+Gt

to combine the concentration of hematite and goethite. The RelHm+Gt record correlates well with percentage content of eolian fractions (Sun & Liu, 2007) from the late Pliocene to early Pleistocene (Figure 7), which further demonstrates that hematite and goethite are weathering products like other silicate fractions and are derived from Asian inland dust sources rather than having volcanogenic or authigenic origins. The DRS proxy succeeds in extracting an eolian signal from ash layers because of insignificant amounts of hematite and goethite in these layers. Compared with chemical separation procedures, DRS measurements are more

rapid and nondestructive. Thus, RelHm+Gt obtained from DRS is a reliable and useful proxy for eolian dust. It is expected that hematite/goethite concentration at Hole 885A may provide detailed information about climate change and the environment of dust sources. Hematite and goethite formation is prevalent under

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opposite conditions, where goethite occurs predominantly in cool and humid climates, while hematite for- mation is favored in warmer and drier climates (Kämpf & Schwertmann, 1983; Schwertmann, 1993). The major Asian dust sources, Taklimakan Desert (Zhang et al., 1993, 1997), Gobi Desert (Sun et al., 2001), and Qaidam Basin (Pullen et al., 2011), have different altitudes and topographic positions, with temperature and precipitation differences among these regions. Hematite/goethite concentrations and their relative proportions may vary with environmental conditions in source regions. The North Pacific Ocean is an extensive sedimentary sink that receives wind-borne fractions from different Asian dust sources (i.e., multiple-sources-one-sink). Thus, on a hemisphere scale, it is unreasonable to expect antiferromagnetic minerals to reflect details of changes in climate and environments of dust sources. Despite all of this, the combination of hematite and goethite as the major components in eolian fractions is capable of identifying eolian signal in pelagic sediments.

5. Conclusions 1. Four types of ferrimagnetic minerals have been identified in the studied sediments from ODP Hole 885A. The first is biogenic magnetite with low coercivity and negligible magnetostatic interactions, which is the dominant magnetic carrier in Unit 1 (before 2.58 Ma) and exists in all samples. The second is SD (titano) magnetite inclusions embedded in silicate hosts with stronger magnetic interactions and higher coerciv- ity. The third type is submicron and subangular vortex state magnetite that mainly occurs after 2.58 Ma. The fourth type is ultrafine SP maghemite, which occurs as polycrystalline aggregates with subangular magnetite profiles that formed through oxidation of coarse magnetite. 2. Changes in ferrimagnetic minerals could reflect increased eolian dust inputs since the late Pliocene. However, traditional environmental magnetic parameters can be affected severely by biogenic and volca- nogenic magnetic components and limit extraction of reliable eolian signals. 3. Volcanic activity generates abundant strongly magnetic minerals and silicates. Volcanogenic magnetic particles are characterized by stronger magnetizations and higher coercivity but contribute little to antiferromagnetic components. 4. Compared with ferrimagnetic minerals, we find that antiferromagnetic minerals are more suitable for indi- cating eolian dust activity in the studied area. Hematite/goethite and silicate fractions deposited at Hole

885A were delivered to the North Pacific Ocean from Asian dust sources. The parameter RelHm+Gt obtained from DRS measurements correlates remarkably well with the percentage content of eolian fractions and is a more reliable and useful eolian dust proxy than magnetic parameters.

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