Controlling the Iron Migration Mechanism for the Cretaceous Sediment Color Variations in Sichuan Basin, China.

Cretaceous continental sediments in Sichuan Basin, China, have different colors, and the reasons for their formation are not determined. Based on mineralogical and geochemical characteristics, red beds and nonred beds in the Upper and Lower Cretaceous sedimentary strata in the western Sichuan Basin are described and tested in this study. The test and analysis of the mineral composition, element content, and iron speciation of mudstone samples with gray-green, gray, and red colors in Cangxi, Bailong, and Guankou formations found that the change in hematite content directly causes the color difference of samples. For red mudstone, the average chemical index of alteration, chemical index of weathering, weather eluviation index (Ba), Ca/(Mg*Al), and Al2O3/SiO2 index are 67.75, 79.94, 2.07, 0.26, and 0.26, respectively, indicating that chemical weathering is the most intense. The geochemical indexes corresponding to gray samples are 64.41, 74.91, 2.08, 0.19, and 0.24, respectively. Those corresponding to the gray-green samples are 62.30, 70.68, 2.17, 0.21, and 0.24, with the weakest chemical weathering. The ratio of Cu/Zn and the enrichment factor of V show that red and nonred bed samples are formed in weak oxidation and weak reduction environments, respectively. The red sample contains the highest content of hematite iron. The gray-green sample mainly represents paramagnetic ferrous in clay minerals. The geochemical contents of the gray sample's three iron elements are slightly different, mainly trivalent iron. The change in iron speciation content in different color samples shows that the Fe element forming hematite in red bed samples may come from the weathering of source rock and clay minerals subjected to secondary weathering. At present, it is confirmed that different colors of samples are related to different weathering degrees of source rocks, which can be related to hot, dry/humid climates. It is necessary to distinguish the climate type in combination with other indicators.

Introduction

Mineral composition and elemental composition of sediments are closely related to the parent rock and paleoclimate. They are usually used to reconstruct the paleoclimate and reveal the intensity of chemical weathering, which plays a vital role in indicating the controlling factors of rock weathering and the environment.[1−3] The iron element widely present in sedimentary rocks is susceptible to changes in redox conditions and is very useful for determining the redox conditions of sediments and rocks.[4,5] At the same time, iron-bearing minerals can make the rock show a specific color. Sediment color is considered one of the best indicators of climate.[6−10] It is used to identify rock types and divide and compare the strata, and it is an important index of climate and environmental change.[11,12] According to the genetic sediments, the color of the sediments can be categorized into inherited color, authigenic color, and secondary color. Both hereditary color and self-generated color are primary colors.[13]

The color of clastic rocks is mainly caused by dyeing substances such as iron-containing compounds and free carbon, and the color can be divided into two broad categories: “red” and “gray”.[11] Red represents the oxidizing environment, and gray represents the reducing environment;[10] yellow to orange, brown, maroon, deep purple, and red belong to the “red” category due to the rock’s iron oxide or hydroxide staining.[14] Dark gray to light gray, brown, bluish-gray, and green belong to the “gray” category, and the color usually darkens as the content of organic carbon or dispersed iron sulfide increases.[11] Terrestrial sediments in the Sichuan Basin contain all types of colors. The red sediments produced in this region are widely used as paleoclimate indicators and for paleomagnetism research studies and geochemistry tests.[15,16] Many researchers believe that sediments of different colors represent the paleoclimatic conditions of hot, dry, or cold.[7,17] These color changes are related to fluctuations in redox conditions before or after deposition.[18] For a long time, the most mainstream view is that the oxidation environment formed under hot and arid climate conditions leads to the formation of terrestrial red beds, and the formation of its red color is related to the existence of disseminated hematite.[19−21] However, researchers have demonstrated that in modern desert environment sediments, the red bed is not typical and, under different climatic conditions such as cold and humid, also can form red deposits.[1,22−24] It has been confirmed that the formation of red color is related to disseminated hematite; there are still controversies about its origin and coloring principle. In recent years, some researchers believe that the color of sediments results from the interaction between the regional structure and the global climate cycle[24,25] and propose the thermal origin of terrestrial red beds.[26] These disputes have led researchers to question whether only considering the color change of sediments can be used as the basis of paleoclimate variations.

The red beds in China primarily formed in the Mesozoic and Cenozoic, especially in the Cretaceous, and the lithology covered the conglomerate to mudstone.[26] Cretaceous terrestrial red beds (Figure ) are widespread in the south, west, and northwest of the Sichuan Basin. According to the sedimentary environment, sedimentary facies, biological facies, and stratigraphic integrity, they are geographically subdivided into four parts: Jiange, Zitong-Bazhong, Yibin, and Chengdu-Ya’an (by the regional geology of the Sichuan province).[27,28] The relatively continuous continental sedimentary strata have been discovered in the northwestern sections of the Sichuan Basin. However, Cretaceous iron speciation composition and color origin in this typical sedimentary basin have not been systematically studied.

Figure 1

Geological map and lithostratigraphy of the Sichuan Basin. (a) Geological sketch map of the Sichuan Basin showing the sampling position. (b) Location of the Sichuan Basin in China. (c) Correlation of the Cretaceous strata in the northwestern and western Sichuan Basin. Reprinted from ref (33). Copyright 2021 American Chemical Society.

Herein, we analyzed and integrated the sample from the Lower and Upper Cretaceous in the Chengdu-Ya’an section and a drillcore through the Chengqiangyan Group in the northern Sichuan Basin. Furthermore, it established the relationship between differences in the degree of weathering of samples and the color of stratum change caused by the regional climate, identified the occurrence forms and types of iron elements in different color rocks, and discussed the iron elements’ migration and occurrence forms in different color formations.

Geological Setting

The Sichuan Basin is located on the western margin of the South China block[27] (Figure a), separated from the Songpan-Ganzi terrane by the Longmenshan thrust belt to the west, from the Qinling Orogen and the Qiyaoshan fold-thrust belt to the southeast, and the Daliangshan fold belt to the south (Figure b).[28,29] These structural belts developed at different periods and in considerably distinct ways control the scale and shape of the Sichuan Basin. In the Sichuan Basin, an intracontinental basin, with an area of 26.0 × 104 km2, the Cretaceous terrestrial red deposits are up to 3000 m thick and cover nearly 40% of the basin area.[27,30]

The sedimentary area began to shrink from the Early Cretaceous, mainly distributed north and west of the Sichuan Basin (Figure b). The Chengqiangyan (CQY) Group is composed of depocenters mainly distributed along the Longmen Mountains thrust belt. The thickest early Cretaceous sediments are Guangyuan and Bazhong in the north of the Sichuan Basin. Previous researchers divided the sequence into Cangxi, Bailong (Figure a), Qiqusi, and Gudian formations from bottom to top. The Cangxi Formation is composed of red sandstone interlayered with red siltstone and thin mudstone, which is unconformably atop the late Jurassic Penglaizhen Formation. The Bailong Formation above the Cangxi Formation is a composition of purplish-red to light purple siltstone and mudstone intercalated with green to light gray terrestrial clastic deposits. The overlying Qiqusi Formation is characterized by red siltstone, sandstone, and mudstone, interbedded with thin grayish-green mudstone. The uppermost Gudian Formation is also marked by red siltstone and mudstone, lighter than those underlying strata (Bureau of Geology and Mineral Resources of Sichuan Province, 1991). The Chengqiangyan Group was initially determined to belong to the late Jurassic based on the stratigraphic correlation, which was likely confused with the Upper Jurassic Penglaizhen Formation. Some calcareous microfossils combined with paleomagnetic data and biostratigraphy determined the Chengqiangyan Group as Early Cretaceous.[27,31,32] Recent studies show that the spore and pollen fossils from the samples provided that the Chengqiangyan Group was tentatively assigned to belong to the Valanginian-to-Hauterivian age.[31]

Figure 2

(a) Stratigraphy and sedimentary logs of the cored section ZK-02 and (b) schematic lithostratigraphic column of the Ya’an area. (c, d) Photographs of the representative portions of the Guankou Formation, Upper Cretaceous. (e, f) Bailong Formation characteristics of red mudstone, sand mudstone, and gray silty mudstone (f) taken at 60 and 190 m from drillcore ZK-02, Lower Cretaceous. (g, h) Cangxi Formation characteristics of red mudstone, silty mudstone, and gray and grayish-green silty mudstone (h) taken at 325 and 425 m from drillcore ZK-02.

The Cretaceous strata are relatively complete in the Chengdu-Ya’an area of the western Sichuan Basin, subdivided into the Tianmashan Formation, Jiaguan Formation, and Guankou Formation (Figure b). A suite of brownish-red, purple, and red clastic deposits characterizes the Tianmashan Formation, unconformably overlying the Upper Jurassic Penglaizhen Formation. At the bottom of the Jiaguan Formation is a purplish-red conglomerate, turned to brown-red and purplish-red siltstone with mudstone in the middle-upper part, which conforms with the overlying Guankou Formation. The Guankou Formation conformably overlies the Jiaguan Formation (Figure c), dominated by brownish-red, purple, and red mudstone and siltstone intercalated with marlite consisting of conglomerate with variable thicknesses at its base. Based on the paleomagnetic study of the Cretaceous strata in Western Sichuan, the top of the Tianmashan Formation represented the base of Barremian. The Jiaguan Formation belongs to Aptian to Santonian, because the bottom of Jiaguan Formation represented the basement of Barremian, while the upper Guankou Formation corresponds to a Campanian–Maastrichtian age (Bureau of Geology and Mineral Resources of Sichuan Province, 1991). Researchers in recent years compiled and modified the literature on the Cretaceous stratigraphic correlation in the Sichuan Basin;[27] the age of Guankou Formation (89.8–66 Ma), Jiaguan Formation (125–89.8 Ma), and Tianmashan Formation (145–125 Ma) is precisely defined. The age of the Tianmashan Formation almost overlaps with that of the Chengqiangyan Group.

Results

Mineralogical Features of Samples

The mineralogy of the bulk rocks was analyzed and showed a similar mineralogical composition (Figure ). Quartz and calcite are the dominant minerals (the average content of quartz is 44.3%, and that of calcite is 13.4%). However, the mineral composition varies among the different colors and ages. There is no noticeable difference in hematite content between the two periods for red samples (the average content in the Late Cretaceous is 2.4%, and that in Early Cretaceous is 2.0%), and the content of hematite content in the grayish-green and black gray samples (average content in the grayish-green sample is 0.8%, and that in the gray sample is 1.6%) is much lower than in red (2.0% on average, Figure e). There are apparent differences in hematite content in samples of different colors.

Figure 3

X-ray diffraction pattern of the typical samples. (a) Lower Cretaceous grayish-green mudstone; (b) Lower Cretaceous red mudstone; (c) Lower Cretaceous gray mudstone; (d) Upper Cretaceous red mudstone. (e) The hematite content of different color samples was measured by XRD, and the color in the figure represents the color of the sample. Abbreviations refer to minerals: Q, quartz; C, calcite; F, feldspar; K, kaolinite; I-S, illite and/or smectite; Chl, chlorite; Anh, anhydrite; Dol, dolomite; Hem, hematite.

Scanning electron microscopy observation combined with energy spectrum analysis shows that the red bed samples have strong mineral aggregation, minor massive minerals are developed and have smooth edges and good abrasiveness, and the inner color of mudstone is consistent. Hematite is cemented by feldspar, quartz, or calcite, in which there are two forms of hematite: clastic hematite with good crystallization and fine granular hematite with poor crystallization. Fine-grained hematite with poor crystallization is associated with clay minerals (such as montmorillonite), which is supposed to be secondary hematite. The relatively well-crystallized hematite has poor roundness, which is supposed to be primary hematite (Figure ).

Figure 4

(a–c) Photomicrographs of hematite with two different crystal forms and (d) X-ray EDS spectrogram. The symbol “+” is the test location; (1), (2), and (3) represent different stages of hematite, and (1) may represent hematite formed by Fe that migrated from clay minerals.

The Whole-Rock Geochemical Features

Among the major elements in all samples, the content of SiO2 is the highest, with an average value of 52.81 wt %. The contents of Al2O3, CaO, and Fe2O3 are 13.67, 8.03, and 5.47 wt %, respectively. The contents of MgO, K2O, and Na2O are 2.92, 2.83, and 1.15 wt %, respectively. These elements account for 86 wt % of the total amount, which is consistent with the characteristics of mineral composition. The total Fe2O3 (TFe2O3) in the Early Cretaceous is higher than that in the Late Cretaceous, which is 5.67 and 5.28 wt % on average. The TFe2O3 content in the Early Cretaceous nonred samples (5.42 wt % on average) is slightly lower than that in the red samples (6.02 wt % on average).

The bulk compositions of the representative samples were used to calculate the geochemical indices for reconstructing paleoclimate conditions and evaluating chemical weathering processes in samples.[3] The CIA index shows that the average value of Guankou Formation in the Late Cretaceous is 67, ranging from 60 to 72; the average value of the Early Cretaceous is 66, ranging from 59 to 72. The Early Cretaceous nonred bed samples are slightly lower than red bed samples (the average value of the grayish-green sample is 61, the red sample is 68, and the garnish yellow sample is 70). All of the values indicate that weak weathering occurred in all types of samples.[34] The CIW index and CIA index are similar; there is little difference between the Early Cretaceous and the Late Cretaceous (78.5 and 79.2, respectively, on average). However, it is worth noting that in the Early Cretaceous, the average CIW index of red mudstone is 79.9, the gray-green sample is 70.1, and the gray sample is 74.91. The sample average value of the Ba index in the Late Cretaceous is 2.54 and that in the Early Cretaceous is 1.68, which is slightly lower than that in the Late Cretaceous. Nonred samples (1.99 on average) are higher than red samples (1.49 on average). The Al/Si ratio values in the Late Cretaceous vary from 0.20 to 0.32; those in the Early Cretaceous vary from 0.18 to 0.35. Notably, the values of the Early Cretaceous grayish-green rock samples are lower than those of the adjacent red samples, which indicates that the red samples experienced more intense weathering when they were formed. The CaO/(MgO × A12O3) results show that the average value of Late Cretaceous samples is 0.37, and that of Early Cretaceous samples is 0.14. From Early to Late Cretaceous, the value of samples gradually increases, and there is no significant difference between the red bed samples and the nonred bed samples. Both indicators show significant differences in the Early Cretaceous and Late Cretaceous.

The ratio of trace elements Cu and Zn reflects the redox state in environmental changes. The Cu/Zn ratios of the Upper Cretaceous and Lower Cretaceous are consistent, both of which are 0.3 on average. However, the samples with different colors of the Lower Cretaceous are slightly different. The red sample ratio is 0.32, the gray sample ratio is 0.29, and the grayish-green sample ratio is 0.23. Trace-metal enrichment factors are shown to be generally superior to bimetal ratios as redox proxies.[35] The value of VEF is consistent with the change in the ratio of Cu and Zn. There is little difference between the Late Cretaceous samples and the Early Cretaceous samples, but the difference between the samples of different colors is more pronounced. The average value of the red sample is 1.04, that of the gray sample is 1.18, and that of the gray-green sample is 1.37.

Systemic Variations of Iron Speciation

Fresh and nonpolluting rock samples with different colors were selected for Mössbauer spectroscopy. The pyrolysis experiments are shown in Figure , and their parameters are summarized in Table . Three absorption peaks appeared in the spectra, including two doublets (D1 and D2) and one sextet. As shown in Figure , the doublet D1 with a minor quadrupole splitting (QS = 0.42–0.88 mm/s) and a minor isomer shift (IS = 0.30–0.43 mm/s) was ascribed to either paramagnetic high-spin ferric iron (para-Fe3+) or iron sulfide.[36] Previous studies have shown that the para-Fe3+ presumably originated from clay minerals, and combined with test parameters, it may represent the ferric iron in smectite.[37,38] The doublets with larger quadrupole splitting (QS = 2.63–2.69 mm/s) and isomer shift (IS = 1.13–1.18 mm/s) were attributed to relatively paramagnetic high-spin ferrous iron (para-Fe2+). According to these Mössbauer parameters (Figure ), the D2 doublet may represent the ferrous iron in clay minerals and maybe the iron in chlorite.[39−41] The sextet with a magnetic hyperfine field (Hi) of approximately 51.0 T to 51.7 T was attributed to hematite (mag-Fe3+), and only hematite can produce magnetic splitting at room temperature.[42,43] There is no significant difference in iron speciation between Early Cretaceous and Late Cretaceous red samples.

Figure 5

Mössbauer spectra of pyrolysis experiment samples analyzed at room temperature (293 K). D1 doublet for ferric iron (para-Fe3+) in ferric hydroxide or clay minerals; D2 doublet for ferrous iron (para-Fe2+) in clay minerals; sextet for magnetic iron in hematite (mag-Fe3+).

Table 2

Values of the Hyperfine Parameters from the Best Fits of 57Fe Mössbauer Spectra for the Samples at Room Temperaturea

name	iron species	relative content %	IS/mm s–1	QS/mm s–1	HW/mm s–1	Hi/T
GJS-06	para-Fe2+	19	1.15 ± 0.01	2.66 ± 0.02	0.34 ± 0.02
	para-Fe3+	14	0.41 ± 0.04	0.55 ± 0.06	0.70 ± 0.13
	mag-Fe3+	67	0.36 ± 0.01	–0.24 ± 0.01	0.32 ± 0.02	51.3 ± 0.0
GJS15	para-Fe2+	29	1.15 ± 0.02	2.69 ± 0.03	0.27 ± 0.04
	para-Fe3+	21	0.35 ± 0.05	0.59 ± 0.06	0.50 ± 0.13
	mag-Fe3+	50	0.34 ± 0.01	–0.24 ± 0.03	0.16 ± 0.05	51.4 ± 0.1
GJS-23	para-Fe2+	14	1.13 ± 0.01	2.66 ± 0.03	0.37 ± 0.04
	para-Fe3+	20	0.35 ± 0.01	0.67 ± 0.04	0.69 ± 0.08
	mag-Fe3+	66	0.37 ± 0.01	–0.23 ± 0.02	0.35 ± 0.02	51.0 ± 0.1
HL-18	para-Fe2+	16	1.15 ± 0.02	2.67 ± 0.05	0.34 ± 0.06
	para-Fe3+	6	0.39 ± 0.22	0.98 ± 0.33	0.66 ± 0.83
	mag-Fe3+	78	0.35 ± 0.03	–0.26 ± 0.02	0.32 ± 0.03	51.4 ± 0.1
HL36	para-Fe2+	25	1.18 ± 0.01	2.68 ± 0.03	0.39 ± 0.04
	para-Fe3+	9	0.43 ± 0.03	0.42 ± 0.05	0.30 ± 0.00
	mag-Fe3+	66	0.41 ± 0.02	–0.23 ± 0.04	0.41 ± 0.05	51.6 ± 0.1
XFP17	para-Fe2+	n.d.	n.d.	n.d.	n.d.
	para-Fe3+	27	0.30 ± 0.03	0.71 ± 0.05	0.56 ± 0.10
	mag-Fe3+	73	0.35 ± 0.05	–0.26 ± 0.03	0.30 ± 0.05	51.6 ± 0.1
WJS01	para-Fe2+	17	1.18 ± 0.05	2.63 ± 0.08	0.46 ± 0.07
	para-Fe3+	10	0.35 ± 0.13	0.88 ± 0.21	0.68 ± 0.19
	mag-Fe3+	73	0.38 ± 0.02	–0.25 ± 0.03	0.42 ± 0.04	51.5 ± 0.1
Z21	para-Fe2+	27	1.16 ± 0.01	2.68 ± 0.01	0.40 ± 0.02
	para-Fe3+	31	0.36 ± 0.01	0.60 ± 0.02	0.62 ± 0.03
	mag-Fe3+	42	0.36 ± 0.02	–0.26 ± 0.02	0.36 ± 0.04	51.7 ± 0.1
Z25	para-Fe2+	27	1.14 ± 0.02	2.69 ± 0.03	0.38 ± 0.05
	para-Fe3+	36	0.39 ± 0.03	0.66 ± 0.04	0.57 ± 0.07
	mag-Fe3+	37	0.33 ± 0.01	–0.18 ± 0.03	0.19 ± 0.04	51.6 ± 0.1
Z34	para-Fe2+	15	1.13 ± 0.01	2.68 ± 0.03	0.39 ± 0.04
	para-Fe3+	26	0.34 ± 0.01	0.64 ± 0.02	0.52 ± 0.04
	mag-Fe3+	59	0.35 ± 0.01	–0.26 ± 0.02	0.35 ± 0.03	51.1 ± 0.1
Z37	para-Fe2+	60	1.16 ± 0.01	2.65 ± 0.02	0.36 ± 0.03
	para-Fe3+	40	0.32 ± 0.04	0.49 ± 0.06	0.61 ± 0.13
	mag-Fe3+	n.d.	n.d.	n.d.	n.d.	n.d.

Note that IS is the isomer shift (relative to α-Fe at RT), QS is the quadrupole splitting, HW is the half width at half maximum, Hi is the hyperfine magnetic field, and relative content is the relative spectral absorption area for each species.

Figure shows the distribution of each Fe species and variation of total Fe contents in all samples studied. It is generally observed that the red and gray mudstone samples are mainly oxidized ferric iron (70–100%), while the green mudstone samples are mainly reduced ferrous iron (60%). The relatively high content of mag-Fe3+ characterizes red mudstone, indicating that iron in the sedimentary rock exists in the form of Fe2O3. At the same time, ferric iron in hematite has an overall upward trend from the Early Cretaceous to the Late Cretaceous.

Figure 6

Contents of iron species in different color samples (samples Z21/25/34/37 are from the Early Cretaceous, and samples WJS-01, XFP-17, HL-18/36, and GJS-06/15/23 are from the Late Cretaceous).

Discussion

The Difference in the Weathering Intensity of Different Color Samples

In the process of chemical weathering, the migration ability of geochemical elements is obviously different. K, Na, Ca, Mg, and other active alkali metals are easily leached out, while other relatively stable elements like Si, Al, and Ti are enriched in the residual phase.[44,45] To eliminate the influence of disturbance factors and magnify the meaning of elements, researchers usually use the sum and ratio of element contents (Figure ). The CIA index reflects the weathering degree of aluminous silicate minerals, especially feldspar, into clay minerals. In the process of chemical weathering of the upper crust, Ca, Na, and K elements will gradually precipitate from feldspar. The higher the CIA value, the stronger the weathering effect on the provenance area.[45−47] However, due to the metasomatism of potassium ions, the content of potassium ions in the sedimentary zones is higher than that of source rocks. Therefore, the CIW index is calculated to exclude the increase in K content caused by K+ metasomatism during the diagenesis process. The higher the CIW value, the more significant the weathering degree of the source area.[48] Cretaceous sediments in the Sichuan Basin show that the values of CIA and CIW are consistent, indicating that diagenesis of sediments is not affected by potassium metasomatism. At the same time, red beds and nonred beds occurred continuously in the Early Cretaceous. Thus, the CIA index can ignore the sedimentary differentiation and recyclization caused by the change of the sedimentary environment.[21,28] Therefore, it can be considered that the CIA index of gray-green beds and the gray beds of nonred sediments is lower than that of red samples, showing that the source area of red samples is subjected to the strongest weathering. Meanwhile, the weathering intensity of the samples in the Late Cretaceous is higher than that in the Early Cretaceous.

Figure 7

Correlation ratio diagram of elements. The color in the figure represents the color of the sample.

The A–CN–K ternary diagram also appraises the weathering alteration. It is used to reflect the chemical weathering trend and the changes in principal components and mineralogy in the weathering process.[45,49] The samples were plotted near the apex of Al2O3 (Figure ), implying negligible potash metasomatism during diagenesis on the studied samples. All samples are almost on the same chemical weathering trend line, and the chemical weathering trend is approximately parallel to the A–CN side. The weathering characteristics of mudstones in this figure indicate that Ca and Na plagioclase are decomposed by the poor stability of the mineral structure. K-feldspar has also been preliminarily decomposed. At the same time, the weathering products of the source rock are mainly illite and montmorillonite, which have not reached the degree of kaolinite yet. For the Early Cretaceous samples, the weathering intensity of grayish-green, gray, and red samples increases in turn, which is consistent with the CIA weathering index. Although the weathering intensity of several samples in the Late Cretaceous is relatively weak, most samples are the same as the red samples in the Early Cretaceous.

Figure 8

A–CN–K ternary plot of samples with different colors.

Using the different geochemical characteristics of elements in the weathering process to calculate the parameter index verifies the above viewpoint. The weathering leaching index (BA) reflects the relationship between active components and inert components; the smaller the ratio, the higher the degree of leaching of active components and the stronger the chemical weathering.[21] Calcium and magnesium are medium or active elements that are dissolved and transported in semi-arid and semi-humid environments.[50] At the same time, the radius of the calcium ion is greater than that of the magnesium ion. CaO/(MgO × Al2O3) can reflect the relative content of authigenic calcium carbonate and indirectly reflect climate change. In Figure , the curves of CaO/MgO and CaO/(MgO × Al2O3) are consistent, indicating that the calcite is mainly authigenic precipitation and the content of terrigenous clastic calcite is minimal. Researchers believe that the geochemical indicators of weathering can usually indirectly reflect the paleoclimate. The Lower Cretaceous ratio changes are generally distributed between high and low values, reflecting the changes and fluctuations of paleoclimatic conditions. In contrast, the proportion of the Upper Cretaceous is relatively high, indicating that the climate of the Late Cretaceous is relatively warm and arid. The paleotemperature of red samples is higher than that of nonred samples. The Cretaceous is generally characterized by alternating arid and semi-arid climate, which belongs to the tropical and subtropical climate context.[51,52] In sedimentary rocks, dolomite is mainly derived from clastic sources.[50] Al2O3 can reflect the change in terrigenous clastic input due to its chemical stability. At the same time, in the process of surface weathering, the aluminosilicate minerals will be transformed into clay minerals. Thence, Al2O3 is inversely proportional to SiO2, showing that silicon to aluminum is proportional to the weathering degree.[53,54] These weathering parameters show the same characteristics as CIA and CIW, which proves that the weathering intensity of nonred bed samples is significantly lower than that of red bed samples. The weathering degree of Late Cretaceous samples is higher than that of Early Cretaceous samples.

Characteristics and Sources of Iron Speciation in Sediments of Different Colors

Previous research studies and the comparison of samples with different colors of the Cretaceous in the Sichuan Basin show that sediment grain size, color, and other characteristics and the change of the deposition environment are closely related.[16,55−57] Compared with marine sediments, terrestrial sediments generally have a higher total iron content and are dominated by ferric iron. Because the chemical properties of trivalent iron are more stable than those of ferrous iron, the weathered iron-bearing minerals in the source area of terrestrial sediments are usually trivalent iron. Iron in continental sediments usually undergoes long-distance migration and transportation.[58,59] Therefore, when the weathering products of the source rocks are transported and rapidly deposited, the highly valent iron is more difficult to migrate. However, when the water body is oxidizing, the highly valent iron complex is more stable, and generally, there is relatively high hematite in the shallow lake sediments. When water is in a reducing environment, especially in a warm and humid environment, high-valence compounds can easily be reduced to low valence by reduction. Therefore, in the deep-water environment, Fe often coexists with much organic matter, resulting in blackening samples.[25,60] In summary, the iron speciation in lacustrine sediments can indicate the primary sedimentary environment.

The morphological characteristics of iron in Sichuan red bed samples show that due to the changes of the sedimentary environment, the reduction environment appears intermittently under the overall characteristics of the oxidation environment. Field and drilling phenomena show that the Cretaceous red bed in the Sichuan Basin occurs in a single layer. The interior of the fresh sample is bright and uniform (Figure a). There is an obvious contact boundary between the upper and lower layers (Figure b). Microscopic characteristics of hematite indicate that it has poor crystallization and is associated with clay minerals (Figure a). There is a linear relationship between Fe and Al contents in red bed samples (Figure a), indicating that the source of Fe in the red bed is mainly a terrigenous input. There is no significant difference in chemical elements such as Al and Si, indicating that the sources of debris in different color samples are similar. Therefore, the poorly crystalline hematite in the red mudstone sample should be authigenic and appear in the sample in a finely dispersed state, similar to the primary hematite formed in the syndepositional-diagenetic early stage in the oceanic red beds.[61] Previous researchers studied a mixture of rhodochrosite, oolitic hematite, rhodochrosite, and oolitic hematite. They found that the crystallinity is inversely proportional to the dyeing ability of hematite. Iron-bearing minerals are essential color factors in rocks: hematite is red, goethite is brownish-yellow, and iron hydroxide is often brownish-red.[62] Combined with the hematite content in the sample, the difference of ferric is the main reason for the color difference of the sample, and the mineral that causes the color difference is hematite.

Figure 9

Field photographs of sample features observed in the representative outcrop. (a) Characteristics of Late Cretaceous red bed samples. (b) Field production characteristics of different color samples in the Early Cretaceous. (c, d) Gypsum minerals widely occurring in the Late Cretaceous red beds of the Sichuan Basin.

Figure 10

Geochemistry and clay mineral characteristics of Early Cretaceous samples from the Sichuan Basin. (a) Correlation between Fe and Al in samples. (b) Variation of iron species with color in Sichuan Basin samples.

Iron in nature exists in the main minerals (mainly silicates) in the form of Fe2+ and will release during weathering in an acidic environment with a low pH. In tetrahedral or octahedral minerals, Fe often exists in layered silicate structures in different divalent and trivalent forms, appears in lamellar clay minerals and hydroxide intercalation in the form of exchangeable cations, adsorbs on the edge of mineral particles, or appears on the surface of clay minerals in the form of iron oxides.[63,64] When the temperature rises, cation exchange occurs, or even if suspended in water, Fe2+ in montmorillonite will be oxidized.[65] The Mössbauer spectrum parameters of the samples measured in the previous show that para-Fe2+ is Fe2+ in chlorite, para-Fe3+ is Fe3+ in montmorillonite, and mag-Fe3+ is Fe3+ in hematite. XRD data have confirmed the above inference. It can be seen from Figure b that with the gradual increase in Fe content in hematite in Sichuan Cretaceous red bed samples, the Fe content in clay minerals gradually decreases. Previous studies show that the composition of clay minerals in the Sichuan Basin showed that the content of montmorillonite is inversely proportional to the content of illite.[66] At the same time, in the early stage of diagenesis or the case of seasonal weathering at surface temperature, when illitization occurs in montmorillonite, the Al element in the environment will displace Fe and Mg ions from the structure and interlayer of montmorillonite.[63,64,67] The alternation of gray and gray-green samples indicates that there may not be a shortage of organic matter at that time. However, in an acidic or organic-rich environment, the Al-montmorillonite can be transformed into kaolinite.[68] In this process, the iron ions that migrated from the montmorillonite enter the water body, combine with other anions, and eventually form hematite in an oxidizing environment or form a divalent iron compound in a reducing environment.

Due to the different oxygen fugacity values of the medium, Cu, Zn, and other copper group elements can be separated during the deposition process. The ratio of Cu/Zn varies with the change in oxygen fugacity of the medium, which is greater than 0.5 in the oxidation environment and less than 0.2 in the reduction environment.[69] In an oxidizing environment, vanadium exists as V5+ in the vanadate and then adsorbs on iron hydroxide, manganese hydroxide, or kaolinite. On the other hand, with an anoxic condition, V will be reduced to V4+ or V3+ and accumulated in the sediments under reducing conditions.[70,71] The VEF and Cu/Zn ratio characteristics show that the sedimentary environment of the samples in this study is a weak oxidation/reduction environment. The gray or grayish-green samples are formed in a relatively reducing environment compared with the red samples.

Therefore, the iron ions forming hematite in the red bed may come from the metamorphic rocks or igneous basement, acid igneous rocks near the basin, and the iron elements that migrated from the clay minerals formed by weathering of the source rocks. Oxygen in the environment oxidizes these elements of Fe, forms amorphous iron hydroxide in water, dehydrates or ages to form goethite and wurtzite, and finally turns into hematite.[4,72] Due to the poor permeability of shale, a relatively closed environment will be formed after diagenesis is completed. Therefore, the above processes mainly occur during weathering and transportation. This situation may indicate that the weathering conditions are more intense when the red beds are formed but weaker when nonred beds such as green beds are formed, so chlorite and other minerals formed by weathering of the source rocks are preserved.

Discussion on the Causes of Different Colored Sediments

Compared with marine red beds, continental red beds formed in sedimentary environments more diversely attracted the interest of researchers. Many previous studies have shown that the hot and dry climate forms the terrestrial red beds. Nevertheless, some researchers believe that red terrestrial sediments can still form under other climatic conditions. Therefore, in recent years, there have been more and more studies on the relationship between sediment color and climate change. Gerhard documented continental red beds like red palaeosols in the humid tropical climate and showed that these are rare and not typical in present-day deserts.[73] At the same time, researchers have confirmed that many red sediments similar to red paleosol were not formed in the tropical climate. Middleton divided continental red beds into primary, secondary, and diagenetic red beds.[74] In the 80s of the past century, Turner had demonstrated red bed coloration due to postdepositional processes based on mineralogical, diagenetic, and grain size data.[22] In addition, it is recorded that red beds are found in arid and humid tropical climates, and the simple existence of hematite minerals is insignificant for any specific climatic interpretation.[22] Walker showed that red is due to the presence of a large number of labile materials, such as various mafic minerals and rock fragments,[75,76] while Myrow explained that even a small amount of precursor organic matter might help form a red bed.[77] Dubiel and Smoot compiled several conditions whereupon red bed formation depends.[78] Jiang et al. proved the thermal origin of red beds in the China mainland by heating black mud to turn it into red.[26] Many researchers have well documented the red color of sediments due to buried diagenesis.[11,16,79,80] Other researchers also believed that the red bed was mainly formed by the erosion and redeposition of the old red bed.[81] However, the problem with this hypothesis is the relative scarcity of modern red alluvial sediments.[74] For a long time, the color and shape of minerals were a guide to parent rock information.[74] Sheldon discussed the modern red deserts in Arizona and Australia and attributed the red bed to the source of these sediments and good drainage conditions.[82] Thus, researchers demonstrated the formation of red color in continental beds due to any of the three major factors: burial diagenesis of sediments, preservation of the inherited red color of source rocks/provenance, and the presence of hot, arid or humid, tropical climate.

Hematite in sedimentary rocks originates from the weathering of iron-bearing silicate in igneous or metamorphic rocks at the basin’s margin. Iron ions form the iron hydroxide, which is dehydrated or aged to hematite.[14,83] The high temperature and short-term precipitation will promote the formation of hematite. The long-term, high-temperature, and dry environment combined with the short-term humid environment is more conducive to the growth of hematite. Because the formation of hematite requires enough water to weather minerals, a high temperature promotes dehydration.[4,84]Figure shows that the content of variation hematite is significantly lower in nonred samples than in the red bed samples. However, the total iron content of the nonred bed samples is similar to that of the red bed samples, indicating that the different hematite contents caused by environmental changes are the main reason for the different colors of the sediments.

The Cretaceous Guankou Formation, Bailong Formation, and Cangxi Formation in the Sichuan Basin are dominated by shallow lacustrine sediments. The provenance area around the Sichuan Basin lacks the provenance area for the formation of large-scale red beds.[28,74] Therefore, red beds cannot be formed mainly by erosion and redeposition of old red beds. It is still controversial that goethite on the surface of clastic minerals or clay minerals will dehydrate and oxidize into hematite with the increase in time and burial temperature because the burial temperature of Paleogene and Neogene red beds has not reached the boundary of transformation.[24] Li reconstructed the mean annual precipitation (MAP) and mean annual temperature (MAT) data of the Early Cretaceous and Late Cretaceous in the Sichuan Basin,[21] indicating that the whole Cretaceous was a long-term semi-arid and arid temperate climate. Previous studies have shown that the climate of the Late Cretaceous was dry and hot through the study of palynology and clay minerals in the Sichuan Basin. In contrast, the climate in the Early Cretaceous was relatively warm and humid.[31,85,86] The climate is consistent with the climatic conditions related to the genesis of terrestrial red beds. Therefore, the Cretaceous red beds in the Sichuan Basin may be due to different climatic conditions, resulting in different sedimentary environments, leading to different degrees of weathering of the source rocks and forming sedimentary rocks of different colors. The Late Cretaceous has a hotter and more arid climate than the Early Cretaceous, and the climate conditions were relatively more stable. As a result, in the Late Cretaceous section, there is no formation of multiple-colored strata like the samples of the Early Cretaceous but a large section of red strata.

In clay minerals, chlorite is usually the product of the relatively physical weathering of parent rock in a cold climate.[87,88] It is often preserved in areas where chemical weathering is inhibited. The existence of chlorite in Early Cretaceous samples and higher characteristic peaks of chlorite in nonred bed samples confirm this view. However, the average temperature of the whole Cretaceous period is high, so the alternating color change of the stratigraphic sequence in the Sichuan Basin may be caused by the intermittent humid environment under the condition of long-term drought and heat. In summary, the periodic color change of Early Cretaceous strata in the Sichuan Basin is related to oxygen fluctuation in the sedimentary environment, which is a climate in which evaporation is more significant than precipitation or precipitation is more remarkable than evaporation alternately. Due to the frequent change of water level and enhanced weathering, the migration of Fe from silicate or clay minerals is accelerated. Therefore, because of the drastic climate change in the Early Cretaceous, the total iron content of red bed samples in the Early Cretaceous is relatively higher than that in the Late Cretaceous. When the environment is depositional where evaporation is more significant than precipitation, the sediment is more likely to be oxidized due to the decrease in water level, forming amorphous hematite; in the opposite case, a relatively stable deep-water environment will be formed. This relatively closed reducing environment leads to the retention of more iron content and weak weathering in the sample. Grayish-green may represent the original color of the sediment.

Conclusions

The main reason for the different colors of Cretaceous samples in the Sichuan Basin is the various hematite contents. The weathering of red samples is the strongest, while gray-green and gray samples are relatively weak. Calcium and sodium plagioclase are decomposed in the red bed samples, and potassium feldspar also has a preliminary decomposition. The iron element forming hematite in the red bed sample may come from the weathered source rock and secondary chemical weathering of clay minerals. Climate change has changed the sedimentary environment, resulting in different weathering degrees of source rocks and eventually forming different colors of sediments.

Materials and Methods

The samples of the Lower Cretaceous were gathered from the ZK-02 drillcore in Wangcang County, the north of Sichuan (Figure b). This core was passed through the Cangxi and Bailong formations, divided by a thick sandstone bed. Red silty mudstone, siltstone, light gray sandstone, and few grayish-green or black gray mudstone jointly constitute the Cangxi and Bailong formations (Figure a). A total of 37 samples were collected and focused on different colors from the core. In the outcrop section of Chengdu-Ya’an in the west of the Sichuan Basin, samples of the Upper Cretaceous were collected by excavating and exposing the fresh outcrop section (Figure b). According to the principle of collecting fresh and uniform mudstone and siltstone along the horizon, the surface weathered the layer and simultaneously removed floating soil. A total of 27 red sandy-mudstone or mudstone samples were collected from this section. All samples were transported to the laboratory in closed iron-free containers for analysis.

All of the samples were crushed into powder using an agate mortar and pestle. Moreover, all samples were stored under dry and hermetic conditions to avoid contamination and minimize chemical variations of the original iron components. X-ray diffraction, X-ray fluorescence spectrometry, and Mössbauer spectroscopy were used to analyze the mineral composition, major elemental contents, and iron speciation of the sample power.

The hematite microstructure and composition characteristics in red bed samples were observed by argon-ion profiling field emission electron microscopy (FE-SEM) and energy dispersive X-ray spectrometry. First, the core of the vertical sample was cut, and then the optical sheet was prepared by a Leica EM TIC 3X three ion beam argon-ion profiler. During the profiling process, 5.5 and 2.0 kV accelerating voltages were selected alternately for four times, with a total of 4 h. The samples after argon ion profiling were treated by conductive metal film deposition, and a Zeiss Sigma field emission electron microscope was used to detect the section surface directly. The area to be measured was delineated for X-ray energy spectrum analysis. Then, an IE250X-Max50 Oxford energy spectrometer was used to analyze and test the mineral composition. The accelerating voltage was 20 kV, the dead time was 35–40%, and the lifetime was 100 s. The energy resolution was 129 eV.

The bulk mineralogical composition was evaluated by X-ray diffraction with Cu-Kα radiation, operated at 40 kV and 40 mA, taking DS (divergence slit) = SS (scattering slit) = 1° and RS (receiving slit) = 0.15 mm. The scanning angle ranged from 2° to52° with a step interval of 0.02° at a rate of 4° (2θ)/min. The software MDI Jade 5 was used to determine the mineral composition of the sample. The percentage of minerals was determined according to the industrial standard of China SY/T 5163-2010.[89] A fully automated sequential wavelength dispersive X-ray fluorescence spectrometer (AXIOS, PANalytical B.V., Netherlands) with Super Sharp Tube of Rh-anode, 4.0 kW, 60 kV, 160 mA, 75 μ UHT Be end Window was used for elemental analysis. About 250–350 mg of powdered sample was gently pressed into a brass sample holder (16 mm in diameter, 1 mm thick) for 57Fe Mössbauer spectroscopy analysis. The brass sample holder was closed at both ends with an iron-free plastic tap. Mössbauer spectroscopy was performed at 293 K using an MA-260 (Bench MB-500) Mössbauer spectrometer with a γ-ray source of 0.925 GBq 57Co/Rh. The measurement and curve-fitting procedures were described elsewhere.[90] The measured spectra were fitted to Lorentzian line shapes using standard line shape fitting routines. The half-width and peak intensities of the quadruple doublet were constrained to be equal. Isomer shifts were expressed concerning the centroid of the spectrum of metallic iron foil.[37]

The chemical weathering for the sediments can be estimated from the chemical index of alteration (CIA). This proxy is calculated from the following equation: CIA = molar (Al2O3)/molar (CaO* + Al2O3 + Na2O + K2O) × 100%;[3,44] CaO* stands for CaO in silicate minerals and is corrected by the method provided by McLennan.[46] The chemical index of weathering (CIW = molar (Al2O3)/molar (CaO* + Al2O3 + Na2O) × 100%),[48] weather eluviation index (Ba = (CaO + K2O + Na2O + MgO)/Al2O3; the oxide index is also the number of molecular moles[91]), indicators for climate change such as CaO/(MgO × A12O3),[92] and the index of clayeyness A12O3/SiO2[53] were also used. The calculation results are shown in Table . Auxiliary estimation of paleo-oxidation–reduction conditions by using the ratio of Cu and Zn and trace element enrichment factor (VEF = (V/Al)sample/(V/Al)ucc) was also conducted.[93]

Table 1

Whole-Rock Geochemical Analysis of the Representative Samples in the Sichuan Basin (wt %)

no.	name	stage	color	SiO2	Al2O3	TFe2O3	CaO	MgO	K2O	Na2O	CIA	CIW	Ba	CaO/(MgO × Al2O3)	CaO/MgO	Al2O3/SiO2	Cu	Zn	V	VEF
1	GJS2	K2g	red	45.17	10.31	5.95	16.59	2.22	1.91	1.46	60.0	68.2	3.91	0.72	7.47	0.23	20.3	58.6	55.9	0.8
2	GJS6	K2g	red	56.59	17.61	6.64	0.61	4.55	3.85	2.36	66.0	69.4	1.17	0.01	0.13	0.31	18.2	137.4	122.5	1.1
3	GJS8	K2g	red	50.92	12.86	4.58	11.35	2.79	2.63	1.26	64.7	75.6	2.54	0.32	4.07	0.25	22.1	71	97.3	1.2
4	GJS11	K2g	red	50.89	12.59	4.97	11.48	2.74	2.55	1.43	62.8	72.9	2.62	0.33	4.18	0.25	23.7	72.6	102.2	1.3
5	GJS14	K2g	red	41.32	12.81	4.72	15.83	4.11	2.84	0.62	71.4	86.2	3.38	0.30	3.85	0.31	20.1	79.1	114.0	1.4
6	GJS15	K2g	red	48.34	12.82	4.77	12.78	2.86	2.72	1.03	67.0	79.1	2.74	0.35	4.48	0.27	19.4	79.3	99.9	1.2
7	GJS17	K2g	red	45.70	10.81	4.36	15.97	2.45	2.19	1.18	63.4	73.6	3.66	0.60	6.51	0.24	20	68.5	68.2	1.0
8	GJS20	K2g	red	36.22	10.95	4.18	22.52	2.43	2.50	0.56	70.7	85.6	4.63	0.85	9.28	0.30	24.2	69.3	80.9	1.2
9	GJS22	K2g	red	42.62	13.69	5.14	15.56	2.79	3.11	0.71	70.6	85.5	2.91	0.41	5.58	0.32	16.5	82.4	107.7	1.2
10	GJS23	K2g	red	46.43	14.83	5.91	12.01	3.61	3.47	0.61	72.0	88.0	2.41	0.22	3.33	0.32	21.9	99.6	119.1	1.3
11	XFP2	K2g	red	55.55	13.57	5.65	7.90	3.28	2.78	0.88	69.7	82.4	2.00	0.18	2.41	0.24	28.3	81.8	107.6	1.2
12	XFP3	K2g	red	52.28	13.13	5.42	10.55	2.83	2.80	0.91	68.5	81.4	2.35	0.28	3.73	0.25	24.2	74	107.4	1.3
13	XFP4	K2g	red	55.03	13.24	5.18	8.83	2.72	2.71	1.05	67.4	79.3	2.09	0.25	3.25	0.24	25.3	68	113.7	1.3
14	XFP5	K2g	red	52.26	12.39	5.41	10.30	2.94	2.40	1.34	63.9	73.7	2.50	0.28	3.51	0.24	30.3	77.3	100.2	1.3
15	XFP6	K2g	red	54.65	14.23	6.10	7.58	3.28	3.05	0.83	70.2	83.8	1.88	0.16	2.31	0.26	29	84.6	116.2	1.3
16	XFP7	K2g	red	49.30	13.62	6.41	10.84	3.55	2.91	0.68	71.7	86.0	2.42	0.22	3.05	0.28	36.8	105.6	108.3	1.2
17	XFP8	K2g	red	52.18	13.06	5.57	10.33	3.02	2.79	0.90	68.6	81.6	2.37	0.26	3.42	0.25	26.6	85.4	109.2	1.3
18	XFP9	K2g	red	43.35	11.35	4.59	18.38	2.71	2.43	0.63	70.8	84.6	3.87	0.60	6.79	0.26	26.4	72.3	102.2	1.4
19	XFP10	K2g	red	53.03	13.12	5.26	9.67	2.90	2.78	0.87	69.0	82.0	2.24	0.25	3.34	0.25	29.2	76.8	101.3	1.2
20	XFP11	K2g	red	55.27	13.46	5.24	8.72	2.36	2.72	0.97	68.7	80.9	1.96	0.27	3.69	0.24	29.2	68.5	102.8	1.2
21	XFP15	K2g	red	46.30	11.44	4.42	15.66	2.03	2.38	0.98	66.4	78.1	3.30	0.68	7.72	0.25	21	59.7	96.9	1.3
22	XFP16	K2g	red	52.86	10.57	5.83	12.86	1.63	2.07	1.28	62.1	71.5	3.01	0.75	7.91	0.20	22.4	47.4	82.4	1.2
23	XFP17	K2g	red	56.69	12.39	4.68	8.54	2.23	2.80	0.97	66.5	79.5	2.08	0.31	3.82	0.22	22.9	57.7	83.8	1.1
24	XFP19	K2g	red	55.04	12.61	4.19	9.96	2.12	2.49	1.07	67.0	78.2	2.21	0.37	4.70	0.23	23.6	56.6	93.8	1.2
25	HL18	K2j	red	63.81	12.99	5.37	6.40	0.86	2.01	1.78	61.8	68.9	1.46	0.57	7.41	0.20	9	18.5	59.6	1.0
26	HL36	K2j	red	53.73	15.37	6.25	6.68	3.56	3.31	1.11	68.0	80.8	1.73	0.12	1.88	0.29	33.7	105.7	119.8	1.2
27	WJS1	K2j	red	50.66	13.25	5.88	4.01	0.77	2.84	0.90	68.7	81.7	1.04	0.39	5.23	0.26	10.3	17.1	41.7	0.7
28	Z1	K1b	yellow-gray	53.77	14.14	5.74	5.76	3.11	3.08	0.87	69.5	83.2	1.63	0.13	1.85	0.26	20.2	81.7	90.1	1.0
29	Z3	K1b	red	52.30	15.23	6.36	5.29	2.95	3.42	1.01	68.5	82.1	1.47	0.12	1.79	0.29	31.7	98.1	97.9	1.0
30	Z4	K1b	grayish-green	59.32	13.68	4.01	5.04	2.47	2.72	1.35	64.9	75.5	1.51	0.15	2.04	0.23	20.6	84.8	102.9	1.2
31	Z5	K1b	red	53.77	13.24	4.84	6.76	2.93	2.58	0.94	69.2	81.0	1.82	0.17	2.31	0.25	23.8	77	86.6	1.0
32	Z6	K1b	gray	54.78	13.54	5.07	5.55	3.80	2.59	1.23	66.4	77.0	1.81	0.11	1.46	0.25	21.2	79.2	97.6	1.1
33	Z8	K1b	grayish-yellow	55.03	19.05	6.56	1.99	2.75	4.49	0.86	71.2	87.0	0.88	0.04	0.72	0.35	38.5	109.9	107.1	0.9
34	Z9	K1b	grayish-green	52.48	14.19	6.17	6.36	3.86	2.69	1.61	63.4	72.8	1.89	0.12	1.65	0.27	31.2	97.3	109.8	1.2
35	Z10	K1b	red	57.93	16.03	7.21	1.37	3.80	3.03	0.92	71.8	84.1	1.05	0.02	0.36	0.28	39.5	111.6	115.8	1.1
36	Z11	K1b	red	56.67	16.22	7.16	2.26	3.23	3.85	0.99	68.6	83.3	1.11	0.04	0.70	0.29	38.1	110.8	100.9	1.0
37	Z12	K1b	red	55.81	14.06	5.10	4.92	2.60	2.90	1.15	67.0	78.8	1.46	0.13	1.89	0.25	33.9	82.4	94.2	1.0
38	Z13	K1b	red	53.43	12.80	5.61	6.84	2.64	2.54	1.27	64.9	75.4	1.87	0.20	2.59	0.24	30.2	82.6	88.7	1.1
39	Z14	K1b	gray	54.49	13.15	5.09	6.56	3.39	2.98	1.05	66.4	79.2	1.93	0.15	1.94	0.24	21.8	80	102.4	1.2
40	Z15	K1b	grayish-green	43.84	12.00	3.12	16.63	2.68	1.91	1.82	59.8	66.7	3.51	0.52	6.22	0.27	20.5	80.5	118.9	1.2
41	Z16	K1b	red	54.68	17.53	6.71	2.58	2.86	4.30	0.80	70.6	86.9	1.02	0.05	0.90	0.32	32.9	109.8	101.7	0.9
42	Z17	K1b	red	59.38	16.90	5.48	1.40	2.87	3.57	1.15	68.9	81.7	0.92	0.03	0.49	0.28	29.9	82.8	94.9	0.9
43	Z18	K1b	red	53.63	17.21	8.02	2.10	3.21	3.78	0.79	72.0	86.9	1.01	0.04	0.65	0.32	37.8	113.3	102.8	0.9
44	Z19	K1b	red	56.21	13.67	5.08	5.12	2.98	2.64	1.25	66.3	76.9	1.59	0.13	1.72	0.24	34	89.2	94.4	1.1
45	Z20	K1b	red	57.75	12.47	4.52	5.53	2.67	2.40	1.18	65.8	76.2	1.71	0.17	2.07	0.22	22	68.6	81.8	1.0
46	Z21	K1c	gray	51.26	14.08	5.86	7.61	3.32	2.99	1.95	59.3	68.7	2.04	0.16	2.30	0.27	27.1	95.4	96.4	1.1
47	Z22	K1c	grayish-green	57.24	10.79	3.90	7.41	3.90	1.11	1.45	64.4	69.4	2.49	0.18	1.90	0.19	6.2	47.1	115.2	1.7
48	Z23	K1c	red	51.36	12.47	5.14	8.84	2.75	2.54	0.96	67.8	79.7	2.19	0.26	3.21	0.24	29.8	80.5	94.2	1.2
49	Z24	K1c	red	56.15	13.91	5.90	4.59	3.22	2.70	1.20	67.0	78.0	1.54	0.10	1.42	0.25	28.6	86.7	92.6	1.0
50	Z25	K1c	red	57.43	16.09	6.94	2.53	3.05	3.43	1.10	68.7	81.6	1.11	0.05	0.83	0.28	33.4	105.4	102.0	1.0
51	Z26	K1c	red	49.37	15.06	7.25	6.42	3.43	3.31	0.92	69.6	83.3	1.69	0.12	1.87	0.31	37.1	112.9	106.0	1.1
52	Z27	K1c	red	52.30	13.40	5.52	7.85	2.78	2.76	0.91	69.1	81.7	1.93	0.21	2.82	0.26	32.9	92.1	100.9	1.2
53	Z28	K1c	red	57.02	15.61	6.55	3.21	3.23	3.22	1.03	69.5	82.2	1.23	0.06	0.99	0.27	28.9	102.7	95.8	1.0
54	Z29	K1c	red	54.39	13.98	5.29	5.08	3.43	2.86	1.14	67.1	78.8	1.64	0.11	1.48	0.26	21.7	83.8	82.9	0.9
55	Z30	K1c	red	54.83	15.83	6.64	4.18	3.05	3.63	1.05	68.2	82.1	1.33	0.09	1.37	0.29	37.3	109.9	98.8	1.0
56	Z31	K1c	red	52.10	14.61	6.00	6.30	3.36	3.06	1.10	67.9	80.2	1.72	0.13	1.87	0.28	32.7	102.5	109.4	1.2
57	Z32	K1c	gray	56.19	10.16	3.43	8.13	2.83	1.61	1.37	61.9	69.3	2.55	0.28	2.87	0.18	14.4	50.1	80.8	1.2
58	Z33	K1c	grayish-green	61.66	11.06	4.21	5.39	2.57	1.74	1.55	61.3	68.5	1.87	0.19	2.09	0.18	13.2	56.2	117.5	1.7
59	Z34	K1c	red	54.62	14.27	6.01	5.10	3.30	2.81	1.24	66.8	77.8	1.59	0.11	1.54	0.26	27	92.9	105.1	1.2
60	Z35	K1c	gray	51.38	12.75	5.39	8.17	2.78	2.65	0.95	68.1	80.3	2.06	0.23	2.94	0.25	28.5	85.7	99.5	1.2
61	Z36	K1c	red	53.19	13.43	5.30	7.08	2.55	2.98	1.05	66.8	79.5	1.81	0.21	2.78	0.25	34.9	87.5	85.6	1.0
62	Z37	K1c	grayish-green	51.97	15.62	5.37	6.24	3.34	3.79	1.92	60.0	71.2	1.73	0.12	1.87	0.30	23.2	110.1	126.9	1.3