Origin of clay minerals on section of Luochuan loess-palaeosol in Shaanxi Province, northwest China

Abstract

Crystallinity, polytype, and morphology of clay minerals in the Luochuan loess-palaeosols in Shaanxi Province, northwest China were studied in order to have an insight into their origin using X-ray diffraction (XRD) analysis and scanning electron microscopy (SEM) methods. The SEM observations showed that the morphology of some illites seems to be lath-shaped crystals. An analysis of illite crystallinity (IC) on the Luochuan section indicated that the origin of illite was related to the rocks of an anchizone. Most illite in the Luochuan loess-palaeosol section were of 2M₁ polytype, but some were of 1M polytype formed by weathering of feldspar in the process of pedogenesis. Illite in the Luochuan section has undergone both physical and chemical weathering. These results revealed that most illite were of detrital origin related to the source area of an anchizone, but parts of the illite were of an authigenic origin formed during pedogenesis after sedimentation. Chlorites in the samples of the Luochuan section were mainly composed of irregular flaky grains and their crystallinity was good. These showed that chlorite had the detrital origin formed by physical weathering. Kaolinite crystallinity was relatively good. The value of CIA on the Luochuan section ranges from 61.9 to 69, and therefore kaolinite could not be formed during weathering and pedogenesis. These results indicated that the kaolinite had a detrital origin. Morphology of smectite seems to be capillaceous. The XRD patterns of all samples contained diffraction peaks at 1.5218 Å (nontronite) and 1.5052 Å (montmorillonite), thus indicating an intermediate composition between trioctahedral and dioctahedral smectite. The smectite crystallinity was very poor. These results revealed that smectite in the study area was authigenic rather than detrital origin.

Keywords

clay minerals / crystallinity / loess-palaeosol / origin / polytype

Introduction

The origin of clay minerals is often intricate, but understanding of origin of clay minerals is required for a paleoclimatic reconstruction. The clay minerals in recent sediments usually contain both detrital and authigenic clays, and the detrital minerals account for the majority of clay minerals (^{Singer, 1984}; Eslinger and Pevear, 1985; ^{Trindade et al., 2013}). The hypothesis that most of the clay minerals in recent sediments are of detrital origin has been supported by many investigations. The paleoclimatic parameters of detrital clay minerals are thus, in general, used to explain a climate change in the source area, and those of the authigenic clay minerals are used to explain the climate condition in the sedimentary area.

Identifying whether the detrital clay minerals are physical or chemical weathering products is also necessary for examination of their origin. Whether the authigenic clay minerals formed during sedimentation or during diagenesis should be distinguished. The authigenic clay minerals formed during sedimentation, in addition, contain two kinds (newly formed clay and transition product of already existing silicate minerals), which reflect the sedimentary environment (climate, water medium conditions, and so on) of the depositional period. Diagenetic clay minerals also include both newly formed clay and transformation product, but they reflect the characteristics of the water medium in rock pores.

In general, the loess-palaeosols have undergone two processes of sedimentation and pedogenic weathering. If the clay minerals in palaeosols were formed by weathering of non-clay parent materials, they would be stable minerals under the climatic conditions. If the clay minerals were derived from the clay parent material prior to deposition, the inherited minerals would be the stable products under the climatic environmental conditions of pedogenesis (^{Hong, 2010}).

Because the Luochuan loess-palaeosol section outcrops well and displays continuous sequences, it has been extensively used for interpretation of the Quaternary paleoclimate changes in the Chinese Loess Plateau (CLP) and Eurasia. Although some researchers (^{Liu, 1985}; ^{Zheng et al., 1985}; ^{Bronger and Heinkele, 1990}; ^{Rutter and Ding, 1993}; ^{Ji et al., 1999}; ^{Gylesjö and Arnold, 2006}) have investigated the origin of clay minerals in CLP, the results showed some differences. In this paper, polytype, morphology, illite crystallinity, and chemical index of clay minerals have been investigated in order to explain the origin of clay minerals in the Luochuan loess-palaeosol sequences.

Geological setting

CLP is situated in the second step of China topography, and it is distributed mainly in Shanxi, Shaanxi, Gansu, Qinghai, Ningxia, and Henan Provinces. Its total area is about 410000 km², and its altitude ranges about 1000 to 1500 m. The surface sediments in CLP consist mostly of eolian deposits of the Quaternary (loess-palaeosol) with a few rock outcrops. The thickness of loess-palaeosol sediments ranges from about 50 to 80 m, with a maximum thickness up to 150–180 m.

The CLP consists of the Holocene palaeosol, Pleistocene Malan loess, Pleistocene Lishi loess, Pleistocene Wucheng loess, and Neogene laterite soil (^{Liu, 1985}; ^{An et al., 1998}; ^{Won et al., 2018}). The Luochuan section is composed of grey-yellow to yellow sandy or silts clay soils (loess) and light brown silty clay soils (palaeosol) (Figs. 1 and 2).

Holocene palaeosol (Q₄) (S₀): It contains the Heilu soil, with evidence of biological disturbance and human activity, and its lower part includes calcic sediments. It is about 1.4 m in thickness.

Pleistocene Malan loess (Q₃) (L₁): The upper and lower parts of Malan loess are light grey-yellow to grey-yellow; the middle portion is light brown, with the lithologic character of silty tosandy soil, loose texture, and vertical joint, and containing carbonate white spots; the lower part contains the mixed red clayey soil. It is about 7.6 m thick.

Pleistocene Lishi loess(Q₂) (S₁–L₁₅): The Lishi loess is composed of yellow, grey-yellow, tawny, brown, and light red-yellow sandy or silty clay soils (loess) and light reddish-brown silty clay soils (palaeosol) with loose and porous texture and calcareous concretions. Its thickness is about 75.5 m.

Pleistocene Wucheng loess (Q₁) (S₁₅): The Wucheng loess contains a good development of calcareous concretion, which appears as upright shape along root pores or fissures. It is about 1.2 m in thickness.

Materials and methods

Sample processing and preparation

Genetic analysis of clay minerals from the Luochuan loess-palaeosol section was made of eight samples taken from a relatively good development of L1, L5, L9, and L15 loesses and their interbedded palaeosols (Fig. 1).

The bulk samples for XRD analysis were processed according to the method described in ^{Andreola et al. (2004)}. The<2 mm of clay particles were separated by means of Stokes’ precipitator method (^{Burt, 2004}). Oriented clay samples were prepared by dropping the clay suspension onto a glass slide and were then dried at the room temperature (^{Won et al., 2018}). To distinguish smectite clays in the samples, oriented clay samples were saturated with ethylene glycol at 65°C for 4 h in an electric oven (^{Rateev et al., 1969}). Bulk powder samples of 1 to 2 g (<65 mm) were prepared to distinguish between dioctahedral and trioctahedral clay minerals, and determine illite polytypes in the layered silicate minerals.

XRD analysis

X-ray diffraction was performed with a Panalytical X’ Pert PRO DY2198 diffractometer, which was operated with Ni-filtered Cu Ka radiation at 40 kV and 35 mA. The slit conditions were: DS=SS=1°, RS=0.3 mm. The XRD profiles were recorded within the range from 3°2q to 65°2q for oriented clay samples, from 22°2q to 36°2q for bulk powder samples, with a scan rate of 4°2q∕min. Step size for d(060) measurement of bulk powder samples is 0.004°2q, and the scanning range is 55°2q to 65°2q (^{Trindade et al., 2013}).

SEM observation

The selected bulk samples for SEM observation were cut to ~0.5 cm in diameter and were then gold-coated (Hong et al., 2007). SEM analysis was undertaken on a JSM-5610 scanning electron microscope. The instrument was equipped with an energy-dispersive spectrometer (EDS) system, which can provide the chemical composition of a microscope zone and favor the determination of mineral particles during SEM observation.

Results

XRD analysis results of oriented clay samples

XRD analysis patterns of oriented clay samples showed 10 Å, 14.2 Å, and 7.17 Å peaks to characterize illite, chlorite, and kaolinite respectively (Fig. 3). The peak at 17 Å that was not clear in an air-dried sample appeared weakly at ethylene-glycol solvated samples, showing the presence of minor smectite. The peaks for qualitative evaluation of the degree of crystallization of chlorite (4.73 Å) and kaolinite (7.17 Å) were also clear.

XRD analysis results of bulk powder samples

XRD analysis patterns of bulk powder samples from the study area showed 2M₁ polytype of illite characteristic peaks at 3.49 Å, 3.20 Å, 2.98 Å, 2.85 Å, and 2.79 Å (^{Moore and Reynolds, 1989}), and also contained 1M polytype of illite characteristic peaks at 3.65 Å and 2.93 Å (Fig. 4). The content ratio between 1M and 2M₁ polytype illites were calculated using a ratio of 2.79 Å / 2.58 Å peak area (^{Maxwell and Hower, 1967}). The peak at 2.79 Å indicates (11\overline{\mathrm{6}}) reflection of 2M₁ polytype illite, the peak at 2.58 Å represents (13\bar{\mathrm{1}}), (116), and (20\bar{\mathrm{2}}) reflections of 2M₁ polytype and (13\bar{\mathrm{1}}), (130) reflections of 1M polytype illite. The result revealed that the content of 1M polytype illite in the palaeosol was slightly higher than that in loess (Table 1). The analysis results of (060) reflection in the study area (Figs. 5 and 6) revealed that relatively strong 1.5393 Å and 1.535 Å diffraction peaks appeared in all the samples, indicating an existence of quartz and trioctahedral illite (biotite). A diffraction peak at 1.5218 Å represents nontronite (trioctahedral smectite). Diffraction peaks at 1.4950 Å, 1.4994 Å, and 1.5052 Å that respectively represent kaolinite, muscovite (dioctahedral illite), and montmorillonite (dioctahedral smectite) appeared with a weak peak intensity. These results revealed that most, but not all, smectites in the study area were trioctahedral ones, and illite was a mixture of dioctahedral and trioctahedral, and all of the chlorites and kaolinites were dioctahedral in their structure.

SEM observation

Under the SEM, the loess-palaeosol sediments exhibit a loose texture. Clay mineral particles in the samples of the study area appear as discrete flakes with angular to irregular morphology. Most of illites in the study area exhibit curved and scalloped outlines (^{Won et al., 2018}), but some seem to be lath-shaped crystals (Fig. 7(a)). Smectite shows fleecy and honeycomb morphologies (Fig. 7(b)).

Discussion

Source of loess materials on Luochuan section

The distribution characteristics of clay minerals in sedimentary area are influenced by the mineral composition in their provenance (^{Ijiri et al., 2018}; ^{Mefire et al., 2018}). Unaltered igneous rocks do not include clay minerals, but they may contain a small amount of mica and chlorite. Low or medium grade of metamorphic rocks (e.g., slate, phyllite, and schist) are generally abundant in mica and chlorite, some hydrothermal altered volcanic rocks and pyroclastic rocks may include smectite (trioctahedral and dioctahedral), chlorite, and chlorite-smectite mixed-layer minerals (^{Eslinger and Pevear, 1985}).

The sediments of the Loess Plateau in China are mainly derived from the deserts in northwest China, which are mainly located in three inland basins of China (Qaidam, Tarim, and Junggar). These basins are surrounded by the Qinghai-Tibet Plateau, the Kunlun, Tianshan, and Qilian mountain chains, and Taishan with an average altitude of 4000–5000 m. These mountains consist mainly of Paleozoic epimetamorphic rocks and some of intermediate acidic magmatic rocks, which serve as the main material source of the northwest deserts (^{Liu, 1985}; ^{Ji et al., 1999}; ^{Fang et al., 2006}; ^{Zhao, 2015}). In these deserts, precipitation is lower than potential evaporation, and especially in the Taklimakan Desert (area of 33.3×10⁴ km²), the worldwide “drought” region, the annual average rainfall is 25–40 mm, while the potential evaporation is 2100 mm. The annual average temperature is 11°C, with average temperature ranging from -9°C to -10°C in January, while the lowest temperature goes down to -20°C in winter (^{Zhao, 2015}).

Many other dry saline lakes and lake sediments in the Quaternary exposed to drought serve also as potential materials for eolian transport. The materials in the Taklimakan Desert, which belongs to an ancient orogenic belt-type, have undergone very weak chemical weathering to form sediments with low maturity. In this desert, the mineral dusts include mainly quartz, feldspar, mica, iron oxide, carbonate, and clay minerals, and the main clay minerals may contain illite, montmorillonite, kaolinite, and chlorite. The contents of each clay component here are 55% for illite, 28% for kaolinite, 15% for chlorite, and 1% for montmorillonite, respectively, while in the Qaidam Basin, they are 60% for illite, 24% for kaolinite, 14% for chlorite, and 1% for montmorillonite (^{Zhao, 2015}).

Origin of clay minerals

Clay minerals are formed and transformed by hydrothermal alteration, weathering, sedimentation, pedogenesis, and diagenesis on the ground surface and in the depths of the Earth’s crust (Mana et al., 2017; ^{Austin et al., 2018}; ^{Cai et al., 2018}). The sediments in the study area did not experience diagenetic stage and are all eolian deposits, and therefore the authigenic clay minerals among them are not those formed by hydrothermal metasomatism or diagenesis. From this analysis, it can be preliminarily estimated that the clay minerals in the study area are clastic origin transported from the source area or authigenic origin formed by weathering in situ.

The clay mineral composition in the study area is mainly illite, with minor chlorite, kaolinite, and smectite or illite-smectite mixed-layer clays (I-S). The relative contents of these four major clay minerals are hardly changed (^{Won et al., 2018}). An assemblage of these clay minerals (except smectite) is very similar to that in the provenance area (deserts in northwest China) and the soil in the adjacent areas (^{Zhao et al., 2001}; ^{Zhao and Sun, 2014}; ^{Zhao, 2015}). This implies that the composition of clay minerals in the study area could be affected by the source region. The proportions of each clay mineral, however, exhibit certain difference between those of the source and study areas. Clay minerals in the source areas have a relatively lower content of chlorite than that of kaolinite, while the study area shows the opposite trend.

This may be because the mineral composition in the source area was changed due to sorting during the transportation of mineral dust, weathering, and erosion in the sedimentary area (^{Garşon et al., 2014}). The loess inherited the mineralogical characteristics in the dry source region, but its mineral composition was changed because its heavier particles were first separated from the dust cloud in the process of their transportation (^{Gylesjö and Arnold, 2006}).

Origin of illite: The illite in sedimentary rocks is generally formed either by weak chemical weathering of aluminum silicate minerals (mica or feldspar, etc.) and mafic minerals under the weak alkaline environment and dry climate conditions or by the erosion of sedimentary rocks (^{Chamley, 1989}; ^{Liu et al., 2005}). Dry-cold climate conditions and weak chemical weathering are favorable for the formation and preservation of illite, so that its presence in the sediments can be used as an indicator of dry-cold climate conditions (^{Zhang et al., 2000}; ^{Varga et al., 2011}).

In the weathered zone, illite can be mainly formed in three processes (^{Ji et al., 1999}). First, new illite is formed by weathering of feldspar. Illite formed through this process exhibits a slight expansibility, low layer charge (<0.75), lath-shaped crystals, and 1M polytype. Second, it is formed through illitization of the smectite during a dry-wet climatic cycle. Under the environment of wet–dry climatic cycling on land surface, smectite can be transformed into illite, which is also an important mechanism of the formation of I-S (^{Środoń and Eberl, 1984}). Third, illite is formed by physical and chemical weathering of muscovite (^{Bronger and Heinkele, 1990}).

The crystallinity of illite (Kübler index) is the main index to distinguish diagenetic zone, anchizone, and epizone (^{Frey, 1987}). The anchizone exists in the low-temperature metamorphic zone between the diagenetic zone and epizone. The epizone exists between the anchizone and the upper zone of the metamorphic domain. The boundary between the anchizone and epizone is IC=0.25°D2q, and that between anchizone and diagenesis is IC=0.42°D2q. The temperature range of the anchizone is generally considered to be 280°C–360°C (^{Ji et al., 1999}). The Kübler index of illite in the study area ranged from 0.491°D2q to 0.255°D2q (^{Won et al., 2018}), most of which were within the Kübler index range in anchizone (0.42°D2q– 0.25°D2q). The crystallinity of illite in the study area is similar to that of the illite in the anchizone, indicating that the origin of illite in the study area is related to the rocks of the anchizone.

Most of illites in the study area were of 2M₁ polytype (the existence of 10 Å (001), 5 Å (002), and 3.34 Å (003) peaks) (Fig. 3). The main polytypes of illite include 1M, 2M₁, and 3T. 1M polytype illite has a monoclinic system and its unit cell includes one structural unit layer and 2M₁ polytype illite has a monoclinic system while its unit cell contains two structural unit layers. In the diagenetic stage at low temperature (<280°C), only 1M polytype illite forms instead of 2M₁ polytype illite. The content of 2M₁ polytype illite increases with a rise of temperature in the anchizone (280°C–360°C), and in the epizone (>360°C), all of illites are of 2M₁ polytype (^{Weaver and Broekstra, 1984}; ^{Niwa et al., 2016}). The crystallinity index and polytype features of illite in the study area thus showed that most of the illite originated under the geological environment of anchizone or similar that at a middle-high temperature (>280°C). On the other hand, the XRD analysis results of bulk powder samples showed that the loess-palaeosol in the study area contains illites of both 1M and 2M₁ polytypes (Figs. 3 and 4), indicating the existence of illite unconnected with anchizone. Therefore, the 2M₁ polytype illite and 1M polytype illite can be considered to be detrital and authigenic origins respectively in our study. The proportion of 2M₁ polytype and 1M polytype illites in the Luochuan loess-palaeosol is shown in Table 1. The SEM observations in samples from the study area showed that the morphology of some illites seems to be lath-shaped crystals (Fig. 7). The existence of 1M polytype illite implies that illite might be generated from the weathering of feldspar, in agreement with the SEM observation (^{Eggleton and Buseck, 1980}).

The chemical index of Al-rich (muscovitic) illite is up to 0.4, representing a strong hydrolysis; the chemical index of Mg- and Fe- rich illite (biotite) is less than 0.15, representing illite experiencing only physical weathering; and the index values between 0.15 and 0.4 represent illite of the intermediate composition (Esquevin, 1969). The chemical index of illite in the Luochuan loess-palaeosol ranges from 0.294 to 0.394, the value of 0.299–0.394 for illite in the palaeosol is greater than that of 0.294–0.334 in the loess (^{Won et al., 2018}). These results show that the illite in the study area has an intermediate chemical composition between Fe- and Mg-rich illite and Al-rich illite, and therefore, experiences physical to weak chemical weathering.

The above results (illite crystallinity and polytype analysis) revealed that most of the illite in the study area had a detrital origin related to the anchizone in the source area, but some illites had an authigenic origin from pedogenesis after deposition (from the presence of 1M polytype illite, SEM observations and illite chemical index).

Origin of chlorite: The chlorite in sedimentary rocks can also be formed by degradation of mica or feldspar in igneous and metamorphic rocks, and mafic minerals, or erosion of sedimentary rocks (^{Chamley, 1989}). Chlorite can transform into smectite with an enhancement of weathering (^{Chen and Wang, 2007}). The main cations of chlorite are silicon, aluminum, iron, and magnesium, and thus chlorite can be formed in alkaline medium and weak chemical weathering environment. Cold and arid regions are favorable for formation and preservation of chlorite (^{Varga et al., 2011}). The physical weathering gives rise to the conservation and concentration of chlorite during the erosion cycle; the chemical weathering causes its decomposition. The chlorite in initial soil is generally of a detrital origin (^{Barnhisel and Bertsch, 1989}). From SEM analysis of the samples in the study area, euhedral chlorite with a coniferous or leaf shape was not found, and they are mainly composed of massive aggregates or flakes (^{Won et al., 2018}). Chlorite crystallinity is generally determined from the sharp 4.7 Å peak in the XRD spectrum. In general, crystallinity of clay mineral is generally related to temperature, buried depth, time, and sedimentary medium, which is highly dependent on the buried depth (^{Xu, 2008}). The chlorite crystallinity of the samples in the study area can be qualitatively evaluated from the XRD pattern (Fig. 3). The sharpness of 4.73 Å peak of chlorite is very fine (where 1=sharp and clear peak, and 3=those with ambiguous shape) (^{Chamley, 1989}), thus showing that chlorite crystallinity is good. This revealed that the chlorite in the study area was formed during diagenesis in the deep buried depth, and transported into the sedimentary area after having been uplifted and exposed to the physical weathering. This interpretation came to the conclusion that chlorite in the study area might have a detrital origin from physical weathering.

Origin of kaolinite: Kaolinite is formed by strong leaching of rock under the conditions of humid climate and acidic medium, and its main cations are Si and Al. It is a decomposition product from silicate minerals (especially feldspar, mica, and pyroxene) under the various physiographical environments, and warm and humid climate is therefore favorable for the formation and preservation of kaolinite. The formation of kaolinite on the regional scale is related to a tropical-subtropical humid climate (^{Chamley, 1989}; ^{Hallam et al., 1991}; ^{Righi et al., 1995}; ^{Ruffell et al., 2002b}). It can also be formed at the early diagenetic stage (Pe-Piper et al., 2005; ^{Piper and Normark, 2009}), but not generally form at the early stage of pedogenesis (^{Wilson, 1999}). Kaolinite in palaeosol is usually considered to be detrital mineral, as the formation of kaolinite by surficial weathering is a common process occurring in tropical climate conditions (^{Aristizábal et al., 2005}). The significant differences between the authigenic kaolinite formed by pedogenesis in recent soil and the detrital kaolinite in sediments are that the authigenic kaolinite has a smaller particle size and poor crystallinity, and also the isomorphous substitution of a small amount of Al by Fe (^{Herbillon et al., 1976}).

In the study area, the content of kaolinite in palaeosol is higher than that in loess, which appears to be the weathering product (authigenic kaolinite) of detrital silicate minerals after deposition (^{Won et al., 2018}). On the other hand, the crystallinity of kaolinite (half width value of 7.1 Å diffraction peak) (^{Bauluz et al., 2014}) was relatively higher (Fig. 3). The value of CIA in the study area ranges from 61.9 to 69 (^{Won et al., 2018}), and the feldspar has not yet reached the formation stage of illite, and therefore kaolinite could not be formed during weathering and pedogenesis. These results revealed that the kaolinite in the study area was mainly the physical weathering products of clastic origin. ^{Bronger and Heinkele (1989}, ¹⁹⁹⁰) also concluded that the kaolinite in the study area was mainly detrital mineral.

Origin of smectite: Authigenic smectite is readily formed in environment of dry and wet climatic cycle, under which the cations can be concentrated in the dry period. It is a product of intermediate chemical weathering, and generally formed in the tropical or subtropical regions with low relief where chemical leaching is not strong (^{Chamley, 1989}; ^{Robert and Kennett, 1994}; ^{Wilson, 1999}; ^{Ruffell et al., 2002a}; ^{Fürsich et al., 2005}). Under warm and humid conditions, sufficient precipitation can promote the dissolution of minerals and ion migration, and arid evaporation causes a salinization of pore water and enrichment of Mg²⁺ and dissolved SiO₂, thus depositing Mg-rich montmorillonite (^{Chamley, 1989}; ^{Worden and Morad, 2003}). ^{Singer (1980)} found that the content of montmorillonite in the soil formed from Galilee basalt (northern Israel) decreased with the increase of precipitation. On the other hand, smectite is a common detrital clay mineral, and its presence plays a very important role in the analysis of paleoclimatic environment (^{Singer, 1984}).

The authigenic and detrital smectites can be identified by the differences in their morphology, crystallinity, chemical composition, REE content, and the characteristics of combined minerals.

Under the SEM, the morphology of the clastic smectite is generally flaky, and that of the authigenic smectite is capillaceous or honeycomb. The SEM observations in samples from the study area showed that the morphology of smectite seems to be fleecy (Fig. 7).

Most of the detrital smectite are dioctahedral and authigenic smectite forms with an intermediate composition between trioctahedral and dioctahedral (^{Iacoviello et al., 2012}; ^{Setti et al., 2014}). The powder XRD patterns of all samples from the study area (Figs. 5 and 6) contained diffraction peaks at 1.5218 Å (nontronite) and 1.5052 Å (montmorillonite), thus representing an intermediate composition between trioctahedral and dioctahedral smectites.

The determination between authigenic and detrital smectite may be also based on the crystallinity of the mineral (^{Singer, 1984}). The crystallinity of detrital smectite is generally poor, while the crystallinity of authigenic smectite is relatively better (^{Setti et al., 2014}). A large number of the DSDP (deep sea drilling project) reports and data of soil mineralogy, however, indicate that such determination is not always reliable. Many montmorillonites in vertisols have a good crystallinity, but authigenic montmorillonites formed by alteration under the marine environment have a poor crystallinity. The crystallinity of smectite is generally expressed by the integral breadth of the 17 Å peak, and its value is less than 1.5 for good crystallinity, 1.5–2 for intermediate crystallinity, and greater than 2 for poor crystallinity (^{Ehrmann, 1998}). In the XRD pattern of samples from the study area, the intensity of the 17 Å peak was weak, and the integral breadth was greater than 2 (Fig. 3), showing that the smectite crystallinity was very poor. These results revealed that smectite in the study area might be authigenic rather than detrital in its origin. The content of smectite in the study area is relatively low (0–8%, with an average content of 3%) compared to other clay species (^{Won et al., 2018}). In case the smectite has a detrital origin, all the loess-palaeosol layers of the sedimentary area will contain smectite. The fact that some of the layers do not contain smectite, however, indicates that the smectite in the study section is not a detrital origin. In addition, the content of smectite in loess is usually higher than that in palaeosol (^{Won et al., 2018}). This is probably due to the leaching of sodium, calcium, and magnesium in the soil layer, since the increase of precipitation during sedimentation of palaeosol has a negative influence on the formation and preservation of smectite (^{Chamley, 1989}; ^{Worden and Morad, 2003}).

There is no illite-smectite mixed-layer clay in the source area, but I-S clay species present in the study area. The existence of the I-S was recognized under the TEM observations in the samples of the study area, as the particle with irregularly interstratified I-S lattice fringes was observed in our previous investigation (^{Won et al., 2018}). Thus, the I-S clay is considered to be an authigenic clay mineral formed by the chemical weathering in the pedogenic process in the depositional area.

Conclusions

The good crystallinity of illite and the existence of 2M₁ polytype illite indicated that illite is mostly detrital origin, inherited from parent material. The chemical index (ICI) of illite, the coexistence of both trioctahedral and dioctahedral illites, the lath-shaped morphology, and the presence of 1M polytype illite imply the formation of a small amount of authigenic illite during chemical weathering.

The value of CIA and good crystallinity showed that the kaolinite has a detrital origin. The irregular flaky feature and good crystallinity of the chlorite indicated that it has a detrital origin.

Very poor crystallinity, fleecy morphology, and paragenesis of trioctahedral and dioctahedral smectites imply that the smectite (or illite-smectite mixed-layer) was an authigenic clay mineral formed in a pedogenic process in the depositional area.

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