Based on thin-section identification, scanning electron microscopy, and X-ray diffraction analysis, fresh whole-rock samples were selected and tested for trace elements, carbon, oxygen, strontium stable isotopes, and rare earth elements. All samples were crushed and inspected by binocular microscope to ensure their reliability, then ground to a 200-mesh particle size using an agate mortar, and sent to the Key Laboratory of Marine Geology and Environment, Test Center, Institute of Oceanography, Chinese Academy of Sciences, for testing and analysis. Trace and rare earth elements were analysed using ICP-OES (Inductively Coupled Plasma-Optical Emission Spectrometer, IRIS Intrepid II XSP, American Thermoelectric Corp.) and ICP-MS (Inductively Coupled Plasma-Mass Spectrometry, ELAN9000, PE Co.). Standard samples included GBW07315, GBW07316, BCR-2 and BHVO-2, and the analysis accuracy error (RSD) was < 5%. Carbon, oxygen, and strontium isotope measurements were carried out on a solid isotope mass spectrometer (MAT262). The blank background of the whole process was about 2–5 × 10−10 g with the error expressed as ± 2σ. The determination result of the NBS standard sample was ± 16 (2σ). The analysis results for the above samples met the requirements of this project.
4.1 Trace elements
In this study, 10 dolomite samples were taken from the 12th and 14th sublayers of the lower lithological member of Dalong Formation (Fig. 3) for trace element analysis (Table 2). The trace element geochemical characteristics of the DSDMM are listed in Table 2.
Table 2
List of trace element contents (mg/kg) and eigenvalues of the DSDMM from Longfeng section
element
|
V
|
Cr
|
Co
|
Ni
|
Cu
|
Zn
|
Ga
|
As
|
Rb
|
Sr
|
Ba
|
Zr
|
sample
|
4-3 N
|
205
|
30.5
|
2.89
|
33.7
|
16.5
|
88.8
|
0.767
|
3.71
|
2.57
|
434
|
11.9
|
10.3
|
4-2 N
|
210
|
29.4
|
2.90
|
33.3
|
13.2
|
41.1
|
0.649
|
3.41
|
2.53
|
451
|
13.1
|
10.5
|
4-1 N
|
356
|
47.7
|
4.05
|
50.6
|
19.4
|
51.0
|
1.41
|
4.94
|
10.6
|
424
|
24.6
|
14.7
|
4-5 D
|
817
|
214
|
7.40
|
127
|
118
|
126
|
6.22
|
20.2
|
23.5
|
469
|
69.0
|
59.6
|
4-4 D
|
1555
|
125
|
3.33
|
79.0
|
44.7
|
67.0
|
3.55
|
17.8
|
14.6
|
167
|
49.4
|
26.5
|
4-Y D
|
1011
|
100
|
8.37
|
139
|
73.8
|
99.6
|
2.78
|
11.1
|
20.3
|
501
|
43.6
|
32.3
|
4-Z D
|
268
|
24.3
|
4.07
|
64.3
|
24.6
|
127
|
0.455
|
2.14
|
1.22
|
889
|
17.1
|
7.64
|
4-3 D
|
248
|
30.4
|
5.59
|
75.2
|
48.7
|
138
|
0.826
|
2.56
|
1.10
|
1387
|
22.9
|
111
|
4-2 D
|
94.2
|
9.84
|
2.27
|
33.7
|
6.75
|
24.9
|
0.209
|
1.54
|
0.948
|
1047
|
7.50
|
4.20
|
4-1 D
|
92.1
|
12.6
|
1.79
|
28.1
|
6.38
|
13.1
|
0.220
|
1.41
|
0.717
|
1378
|
8.81
|
4.86
|
average
|
485.63
|
62.374
|
4.266
|
66.39
|
37.203
|
77.65
|
1.709
|
6.881
|
7.809
|
714.7
|
26.79
|
28.16
|
Clark value
|
135
|
119
|
25
|
75
|
55
|
70
|
—
|
1.3
|
90
|
325
|
390
|
123
|
|
trace elements
|
Nb
|
Ta
|
Pb
|
Nb
|
Mo
|
Cs
|
Th
|
U
|
U/ Th
|
△U
|
V/(V+ Ni)
|
Co/ Ni
|
sample
|
4-3 N
|
0.878
|
0.044
|
0.916
|
0.878
|
18.1
|
0.089
|
0.492
|
5.39
|
10.955
|
1.941
|
0.795
|
0.086
|
4-2 N
|
0.930
|
0.047
|
1.28
|
0.930
|
24.2
|
0.062
|
0.498
|
5.45
|
10.944
|
1.941
|
0.613
|
0.089
|
4-1 N
|
1.66
|
0.118
|
3.54
|
1.66
|
51.8
|
0.347
|
1.57
|
7.21
|
4.592
|
1.865
|
0.876
|
0.08
|
4-5 D
|
3.51
|
0.210
|
6.30
|
3.51
|
107
|
1.61
|
2.42
|
27.7
|
11.45
|
1.943
|
0.865
|
0.058
|
4-4 D
|
2.11
|
0.106
|
6.42
|
2.11
|
118
|
1.19
|
1.42
|
9.04
|
6.366
|
1.980
|
0.952
|
0.042
|
4-Y D
|
2.91
|
0.158
|
6.00
|
2.91
|
135
|
0.798
|
2.07
|
19.8
|
9.565
|
1.933
|
0.879
|
0.061
|
4-Z D
|
0.471
|
0.026
|
0.657
|
0.471
|
13.2
|
0.053
|
0.290
|
6.61
|
22.79
|
1.97
|
0.807
|
0.063
|
4-3 D
|
0.522
|
0.026
|
0.332
|
0.522
|
28.6
|
0.062
|
0.367
|
84.8
|
231.06
|
1.95
|
0.767
|
0.074
|
4-2 D
|
0.246
|
0.016
|
0.186
|
0.246
|
14.8
|
0.031
|
0.165
|
5.54
|
33.58
|
1.98
|
0.737
|
0.067
|
4-1 D
|
0.253
|
0.015
|
0.512
|
0.253
|
11.9
|
0.025
|
0.149
|
5.58
|
37.45
|
1.982
|
0.766
|
0.06
|
average
|
1.349
|
0.077
|
2.614
|
1.349
|
52.26
|
0.427
|
0.944
|
17.71
|
36.89
|
1.945
|
0.806
|
0.068
|
Clark value
|
20
|
0.7
|
12.5
|
20
|
__
|
2.6
|
9.6
|
2.7
|
Characteristic ratio
|
Notes: 1) Element Clark values are quoted from Zhou Yongzhang et al., 1994; 2) △U = Utotal − Th/3; 3) Sampling locations and numbers are the same as for Table 1.
4.1.1 Trace element distribution
The contents of DSDMM trace elements varied greatly. Among the 20 elements counted, emission-type elements (V, Zn, As, Mo, Sr, U, etc.) are enriched types with Clark values higher (or slightly higher) than the crustal value. Most of the rest are depleted types lower than the crustal Clark value, while Rb, Ba, Ta and other elements are more depleted types reflective of continental weathering sources. The above characteristics indicate that the deposition of the DSDMM was greatly influenced by deep crustal elements and the supply of terrigenous materials. The trace element distribution characteristics show that the main material source of the micrite mound is closely related to magmatic-hydrothermal activity in the deep crust, with a mainly endogenous inorganic supply. Combined with the fact that micrite mounds are interlayered in a large set of radiolarian siliceous rocks and have been proven to be of hydrothermal bio-composite deposition origin, it can be inferred that the DSDMM was also the product of endogenous inorganic supply-biological organic aggregation deposition in the deep-water basin environment.
4.1.2 Eigenvalues of trace elements and paleo-oxygen facies analysis
Based on the principles and technical methods of analysing the trace element eigenvalues of paleo-oxygen facies (Wu et al., 1999; Wignall et al., 1994; Algeo, 2004), the paleo-oxygen facies appeared suitable for analysing the sedimentary environment during the formation of the DSDMM. The following causes are shown:
1) Paleo-oxygen facies analysis of U/Th ratios. The U/Th ratios of the 10 samples varied greatly, ranging from 4.592 to 231.06 with an average of 36.89 (Table 2). The evaluation criteria proposed by Wu et al. (1999) indicate a reduction environment with severe hypoxia.
2) Analysis of △U in the paleo-oxygen facies. The calculation formula and discrimination of ΔU still uses the standard of Wu et al., (1999). The variation in ΔU in the 10 samples was very small, ranging from 1.865 to 1.982 with an average of 1.945 (Table 2). This also indicates a reduction environment with severe anoxia.
3) Paleo-oxygen facies analysis of V/(V+Ni) ratios. Using the criterion of Wignall et al. (1994), the 10 samples showed a small range of variation in V/(V+Ni) ratios of 0.613 to 0.952 with an average of 0.806 (Table 2). This also indicates an oxygen-deficient–anoxic-reduction environment.
Together, this analysis confirms that the DSDMM occurred in a reduction environment of low-oxygen–severe anoxia in deep water.
4.1.3 Co/Ni ratio and origin of dolomite
Rocks with a hydrothermal depositional origin tend to be rich in Co and Ni, with Ni being more enriched. Under normal conditions, the Co/Ni ratio of hydrothermal sedimentary rocks is < 1.0 (Algeo, 2004). Therefore, the Co/Ni ratio has become an indicator of the genesis of hydrothermal and non-thermal sedimentary rocks and has a wide range of applications in the genetic analysis of hydrothermal sedimentary rocks.
The Co content of the DSDMM is 1.79–8.37 × 10-6 with an average of 4.266 × 10-6, which is less than the crustal gram average (Zhou et al., 1994) and higher than the background value of normal carbonate rocks. The Ni content is 28.1–127×10-6 with an average of 66.39 × 10-6, which is not only close to the average Clark value of the crust but also much higher than the background value of normal carbonate rocks. The Co/Ni ratio varies from 0.042 to 0.089 with an average of 0.068 (Table 2), which is far lower than the standard upper limit for identifying hydrothermal sedimentary rocks of ≤ 1 (Zhou et al., 1994). Therefore, it indicates that the DSDMM has a hydrothermal deposition origin.
4.2 Carbon, oxygen and strontium isotopes
The reliability and representativeness of all samples involved in carbon, oxygen and strontium stable isotope analysis were evaluated before testing. The evaluation standard adopted two standards: a Sr content not less than 200 × 10−6, as proposed by Derry et al. (1989), and an Mn/Sr ratio not higher than 2–3, as proposed by Kaufman et al. (1993). The Sr abundances of all samples collected from the micrite mound dolomite in this study were above 500 × 10-6 and below 400 × 10-6 and the Mn/Sr ratios were < 1, indicating that the diagenetic alteration of all samples was relatively weak, and the primaeval ocean information was well preserved, providing good reliability and representativeness.
4.2.1 Carbon and oxygen isotopes
The carbon and oxygen isotopes (δ13C and δ18O) of carbonate rocks not only reflect the material sources, paleo-geographic environment and paleo-climatic conditions, but also the water salinity and temperature in the sedimentary environment (Zhang, 1985). For example, δ13C mainly depends on water salinity. Higher water salinity results in greater values and the positive offset is dominant. The δ13C and δ18O compositions of carbonate sediments from different sources are also different (Veizer et al., 1986). The δ13C(PDB) of normal seawater is about 0‰, and the δ18O(PDB) is 1‰ to −4.3‰. The negative offset is generally not more than 3‰, which depends on the carbon cycle and intensity resulting from the paleoclimate. The δ13C derived from mantle is generally heavier and the positive offset is dominant. The δ13C derived from organic matter is generally very light with negative offsets predominating. Although the sources and variations in oxygen isotopes are similar to those of carbon isotopes and are related to water properties, the depositional environmental temperature and salinity, oxygen isotopes are more active than carbon isotopes and are more susceptible to deposition-diagenetic temperature, resulting in higher fractionation strengths and negative offset amplitudes. Therefore, carbon isotopes are more reliable and important than oxygen isotopes for providing paleo-marine information and indicating the genesis and material sources of carbonate rocks (Korte et al., 2005).
In this study, eight dolomites and two dolomitic siliceous rocks were collected from small layers 12 and 14 of the lower member of Dalong Formation (Fig. 3) for carbon and oxygen isotope analyses (Table 3). The results show that the carbon and oxygen isotopic compositions of the DSDMM dolomite have the following two distinct characteristics.
Table 3
Carbon and oxygen isotope data of DSDMM dolomite
small layer
|
sample
|
lithology
|
occurrence
|
δ13C
(PDB) ‰
|
δ18O
(PDB) ‰
|
Inside the 4th layer in Fig.3
|
4-3 N
|
Carbon-rich bioclastic siliceous micro-powder dolomite
|
medium-thick
lamellar
|
-4.9
|
-5.0
|
4-2 N
|
-4.8
|
-4.9
|
4-1 N
|
-5.4
|
-4.6
|
4 small layer mounds
body roof
|
4-5 D
|
Carbon-rich cloud radiolarian siliceous rock
|
lamellar
|
-3.5
|
-9.4
|
4-4 D
|
-5.4
|
-7.0
|
the top of the 4th layer in Fig. 3
|
4- Y D
|
Carbon-rich bioclastic siliceous micro-powder dolomite
|
mound
|
-2.6
|
-6.5
|
4- Z D
|
-2.4
|
-6.9
|
4-3 D
|
-5.5
|
-5.5
|
4-2 D
|
-2.0
|
-7.2
|
4-1 D
|
-7.9
|
-6.7
|
Note: Sampling location and number are the same as Table 1
1) The composition range of carbon and oxygen isotopes varies greatly: δ13C(PDB)‰ = −7.9‰ to −2.0‰ with an average of −4.4‰, while δ18O(PDB)‰ = −9.4‰ to −4.6‰ with an average of −6.4‰. These are slightly lower than the average values of δ13C(PDB)‰ (1.39‰) and δ18O(PDB)‰ (−5.47‰) obtained from simultaneous analysis of mud-microcrystalline limestone and brachiopod shells from the same period in the Upper Yangtze Sea. For instance, compared with the composition of carbon and oxygen isotopes in global Permian seawater (Veizer et al., 1986), the δ13C and δ18O values of DSDMM are lower. According to Wiggins et al. (1993), the in-situ δ18O(PDB) distribution of marine carbonate rocks without alteration in the Late Permian is −2.8–2.2‰, which represents the oxygen isotope baseline of Late Permian marine carbonate rocks. Compared with the dolomite δ18O distribution range (−9.4‰ to −4.6‰), the negative offset amplitude is much larger and, combined with the characteristics of well-preserved original rock structures and material components, this shows that the micrite mound dolomite was not of normal seawater depositional origin, nor was it transformed by strong diagenetic alteration, nor can it be the product of diagenetic fluid metasomatism under high temperature during the burial period. The most reasonable explanation for its origin is that the siliceous micrite mounds formed in a relatively-high-temperature submarine hydrothermal exhalation-sedimentation system. There are many similar examples, including 1) Xiagou, Qingxi Sag in Jiuxi Basin, 2) the Permian Lucaogou Formation in Santanghu Area, Xinjiang, 3) the Lower Cretaceous Baiyinchagan Sag, Erlian Basin, Inner Mongolia, and 4) the dolomite of lacustrine hydrothermal deposition in the Permian Fengcheng Formation in the Wuerhe Sag, in the northwestern margin of the Junggar Basin (Zheng et al., 2003, 2006a, 2018; Liu et al., 2011; Guo et al., 2012).
2) Correlation analysis of δ13C and δ18O shows that the correlation coefficient is only 0.35. The correlation deviation is moderate, which indicates that the changes in δ13C and δ18O values during the formation of the DSDMM were not associated with diagenetic alteration.
Combined with the large set of radiolarian siliceous rocks in the rock surrounding the micrite mound, most of the easily altered opal components still maintain a colloidal structure and the material composition characteristics of the original deposition. This, and the well-preserved structures of opaline radiolarian fossils in micrite mounds (Fig. 4D), proves indirectly or directly that siliceous micritic mounds are undoubtedly the product of submarine hydrothermal-exhalation deposition.
4.2.2 Strontium isotopes
The changing strontium isotopic composition of seawater is controlled by two factors (Zhang, 1985). 1) Terrigenous strontium (i.e., 87Sr formed by the decay of radioactive Rb during continental weathering), is mainly derived from the weathering of silicate parent rocks. A large amount of terrigenous 87Sr injected into the ocean can sharply increase the 87Sr/86Sr ratios of seawater during sea level drops. 2) Mantle-derived strontium, which is dominated by 86Sr, is mainly derived from mantle hydrothermal fluids in the deep crust. A large amount of mantle-derived strontium injected into the ocean will reduce the 87Sr/86Sr ratio of seawater during strong volcanic eruptions, which is the opposite of terrigenous strontium. Therefore, the 87Sr/86Sr ratio can be used to indicate sea-level cycles or volcanic eruption events to establish strontium isotope stratigraphic curves for dating and to indicate the fluid source for carbonate mineral precipitation (Barnaby, 2004), such as the hydrothermal fluids that precipitate dolomite.
According to the strontium isotope analysis results, four dolomite and four normal marine limestone samples were collected from the 4th sublayer (Fig. 3) in the lower lithological member of Dalong Formation (Table 4). Four remarkable characteristics are presented for the strontium isotopes of DSDMM.
1) The 87Sr/86Sr ratios of the four siliceous dolomite samples have a relatively concentrated distribution range of 0.7070223 to 0.7071472 with an average of 0.7071087, indicating that the strontium isotopic composition of DSDMM is preserved better.
Table 4
Strontium isotope analysis of DSDMM and normal marine limestone in the first member of Feixianguan Formation
Strata
|
Sampling location
|
Sample number
|
Occurrence
|
Sample lithology
|
87Sr/86Sr
|
Lowest member of the Lower Triassic Feixianguan Formation
|
14 small layers of Figure 3
|
YJY75
|
Lamellar
|
Purple-red argillaceous microcrystalline limestone
|
0.70714575
|
Layer 13 of Figure 3
|
YJY65
|
0.70726700
|
Upper Permian Dalong Formation
|
Upper lithological section, sublayer 11 in Fig. 3
|
YJY54
|
Thin-medium lamellar
|
dark grey mud-microcrystalline limestone
|
0.70820023
|
Middle lithological section, sublayer 8 in Fig. 3
|
YJY32
|
0.70746219
|
Lower part
lithological section
|
Top of 4 small layers in Fig. 3
|
14-3
|
Mound
|
Carbon-rich bioclastic siliceous micro-powder dolomite
|
0.70714724
|
14-2
|
0.70711918
|
14-1
|
0.70702233
|
Inside the 4th layer of Fig. 3
|
14-0
|
Thick layered
|
0.70714604
|
2) For example, Huang et al. (2008) proposed that the 87Sr/86Sr value of seawater was 0.70714–0.70715 at the PTB boundary in the Zhongliangshan area of Chongqing during the same period. The average 87Sr/86Sr value of seawater at the PTB boundary in Sicily, Italy, was 0.707385 (Korte et al., 2003), who also reported that the average 87Sr/86Sr value of seawater at the global PTB boundary was 0.70715 in 2006 (Korte et al., 2006). The fitted 87Sr/86Sr value of seawater in the same period was estimated as 0.7075 by McArthur et al. (2001) in 2001. In the Sr isotope stratigraphic curve at the intersection of the P-T fitted, the mean 87Sr/86Sr value (0.7071087) of the Dalong Formation siliceous dolomite was the lowest published value of all seawater from the same period, indicating that the micrite mound dolomite was not a normal marine deposition. The negatively offset 87Sr/86Sr ratio indicates that mantle-derived strontium was involved in its formation.
3) The 87Sr/86Sr value of normal marine limestone is 0.70714575–0.70820023 and the average is 0.707519 at the P-T intersection in the Upper Yangtze area, and the 87Sr/86Sr ratio of the DSDMM shows a large negative shift in comparison. The former 87Sr/86Sr values are consistent with the average value published by Korte (2003) and the fitted value proposed by McArthur et al. (2001), which also indicates that mantle-derived strontium was involved in the formation of microcrystalline dolomite.
4) It is very interesting that the 87Sr/86Sr ratios of dolomite samples collected from the lower, middle and upper parts of the siliceous dolomitic micrite mound at the top of the four sublayers are 0.707022 (lower than in seawater of the same period), 0.707119 (closed seawater) and 0.707147 (consistent with seawater of the same period), respectively, indicating a positive evolution trend (Fig. 4B). It is believed that this reflects the DSDMM accumulation process as well as a mixing process of hydrothermal fluid ejected from the seafloor with the normal seawater (i.e., the homogenization of mantle-derived strontium carried by hydrothermal fluid and seawater strontium). This feature is evidence that the dolomite precipitated by hydrothermal fluid mainly came from the deep part of the crust, and can also be used as the siliceous dolomitic micrite mound located near the vent of seafloor hydrothermal fluid, and as a basis for judging the gradual enhancement of strontium isotope homogenization between mantle-derived strontium and seawater strontium in the process of hydrothermal exhalation.
5) The 87Sr/86Sr ratio of the layered, bioclastic, siliceous, micro-powder, dolomite interlayer taken from the interior of the 4th sublayer (0.707146) is completely consistent with the range determined by Huang Sijing et al. (2008, 0.70714–0.70715). Obviously, it is the sediment that the deep hydrothermal fluid ejected from the seafloor and the normal seawater strontium isotopes were completely homogenized and were not the product of direct precipitation from deep hydrothermal fluid ejected from the seafloor, nor the product of diagenetic metasomatism. In the dolomite inside and at the top of the four sublayers, there are rich radiolarian fossils, which further indicates that organisms also participated in the formation of siliceous dolomite during the homogenization of mantle-derived strontium and seawater strontium.
4.3 Rare earth elements
The RRE analysis results and eigenvalues of 10 dolomite samples collected from the DSDMM are summarized in Table 5. The geological significance of each REE eigenvalue and partition pattern is described as follows.
4.3.1 ΣREE eigenvalues and LREE/HREE ratios
Table 5
List of REE contents (mg/kg) and eigenvalues of 10 samples of DSDMM dolomite
Sample
|
La
|
Ce
|
Pr
|
Nd
|
Sm
|
Eu
|
Gd
|
Tb
|
Dy
|
Ho
|
Siliceous dolomitic micrite mound dolomite in the Dalong Formation
|
4-3 N
|
15.0
|
32.3
|
3.85
|
14.6
|
3.82
|
0.337
|
3.81
|
0.699
|
4.22
|
0.883
|
4-2 N
|
2.42
|
3.87
|
0.459
|
1.89
|
0.373
|
0.066
|
0.459
|
0.076
|
0.467
|
0.109
|
4-1 N
|
2.44
|
4.05
|
0.488
|
2.05
|
0.476
|
0.089
|
0.475
|
0.085
|
0.478
|
0.111
|
4-5 D
|
11.9
|
17.6
|
2.21
|
8.52
|
1.97
|
0.446
|
2.33
|
0.403
|
2.42
|
0.584
|
4-4 D
|
6.97
|
13.0
|
1.58
|
6.37
|
1.09
|
0.181
|
0.904
|
0.161
|
0.806
|
0.186
|
4- Y D
|
1.42
|
2.42
|
0.307
|
1.39
|
0.416
|
0.110
|
0.586
|
0.112
|
0.643
|
0.165
|
4- Z D
|
7.62
|
13.0
|
1.62
|
6.14
|
1.36
|
0.265
|
1.43
|
0.254
|
1.39
|
0.357
|
4-3 D
|
4.49
|
7.31
|
0.937
|
4.27
|
1.21
|
0.275
|
1.51
|
0.263
|
1.67
|
0.405
|
4-2 D
|
0.832
|
1.56
|
0.184
|
0.746
|
0.181
|
0.040
|
0.183
|
0.031
|
0.174
|
0.037
|
4-1 D
|
0.765
|
1.12
|
0.143
|
0.564
|
0.115
|
0.024
|
0.121
|
0.022
|
0.133
|
0.036
|
average value
|
5.3857
|
9.623
|
1.1778
|
4.654
|
1.1011
|
0.1833
|
1.1808
|
0.2106
|
1.2401
|
0.2873
|
Dalong Formation siliceous rock
(average of 13 samples)
|
10.32
|
21.48
|
2.45
|
9.27
|
1.74
|
0.326
|
1.66
|
0.266
|
1.46
|
0.33
|
Dalong Formation sedimentary tuff
(average of 27 samples)
|
28.502
|
60.611
|
6.7237
|
25.386
|
5.11222
|
0.7516
|
4.6859
|
0.809
|
4.5304
|
0.9902
|
North American shale
|
41.00
|
83.00
|
10.00
|
38.00
|
7.50
|
1.61
|
6.35
|
1.23
|
5.49
|
1.34
|
|
Sample
|
Er
|
Tm
|
Yb
|
Lu
|
ΣREE
|
LREE/
HREE
|
(La/
Ce)N
|
(La/Yb)N
|
δCe
|
δEu
|
siliceous dolomitic micrite mound dolomite in the Dalong Formation
|
4-3 N
|
0.309
|
0.046
|
0.285
|
0.045
|
11.427
|
5.23
|
1.214
|
0.735
|
0.88
|
0.27
|
4-2 N
|
0.312
|
0.050
|
0.331
|
0.054
|
10.936
|
4.89
|
1.255
|
0.638
|
0.86
|
0.21
|
4-1 N
|
2.29
|
0.381
|
2.47
|
0.361
|
85.021
|
4.62
|
0.951
|
0.52
|
1.01
|
0.13
|
4-5 D
|
1.61
|
0.254
|
1.70
|
0.265
|
52.212
|
4.45
|
1.368
|
0.599
|
0.81
|
0.28
|
4-4 D
|
0.500
|
0.074
|
0.502
|
0.077
|
32.401
|
9.11
|
1.083
|
1.189
|
0.93
|
0.28
|
4-Y D
|
1.04
|
0.179
|
1.20
|
0.194
|
36.049
|
4.96
|
1.185
|
0.544
|
0.87
|
0.27
|
4-Z D
|
0.439
|
0.067
|
0.432
|
0.072
|
8.579
|
2.41
|
1.207
|
0.285
|
0.87
|
0.28
|
4-3 D
|
1.12
|
0.172
|
1.03
|
0.170
|
24.832
|
2.92
|
1.25
|
0.375
|
0.84
|
0.27
|
4-2 D
|
0.119
|
0.015
|
0.106
|
0.019
|
4.227
|
5.18
|
1.111
|
0.557
|
0.94
|
0.32
|
4-1 D
|
0.122
|
0.022
|
0.139
|
0.025
|
3.351
|
4.40
|
1.462
|
0.475
|
0.80
|
0.29
|
average value
|
0.7861
|
0.126
|
0.8195
|
0.1282
|
26.9035
|
4.817
|
1.209
|
0.592
|
0.881
|
0.26
|
Dalong Formation siliceous rock
(average of 13 samples)
|
0.95
|
0.15
|
1.07
|
0.17
|
51.636
|
7.367
|
1.01
|
0.78
|
1.017
|
0.834
|
Dalong Formation sedimentary tuff
(average of 27 samples)
|
2.6998
|
0.4236
|
2.6393
|
0.3959
|
144.24
|
|
|
|
|
|
North American Shale
|
3.75
|
0.63
|
3.51
|
0.91
|
203.41
|
7.232
|
Eigenvalues
|
Note: Sampling locations and numbers are the same as in Table 1
Systematic study has revealed the characteristics of REE abundance in sediments caused by marine hot water and normal seawater deposition all over the world. Fleet et al. (1983) concluded that the ΣREE (total rare earth elements) and LREE/HREE ratios of hot water sediments are low, being characterized by non-enrichment and low differentiation of LREEs and HREEs. The ΣREE and LREE/HREE ratios of normal seawater sediments are relatively high and have the basic characteristics of LREE enrichment and high losses and differentiation of HREEs (Fleet, 1983). Therefore, the ΣREE eigenvalue and LREE/HREE ratio have geological significance in identifying hydrothermal and non-hydrothermal sedimentary rocks (Marching et al., 1982). The range of ΣREE variation in DSDMM is 13.184–84.94 × 10-6 with an average of 26.9 × 10-6. The LREE/HREE ratio is 2.41–9.11 with an average value of 4.82, which is significantly lower than that of North American shale (Table 5), indicating that DSDMM has the genetic characteristics of hydrothermal deposition.
4.3.2 (La/Ce) N ratio
This ratio is an eigenvalue describing the degree of LREE differentiation in rocks. The degree of LREE differentiation in different sedimentary basins can be used to judge their nature. For example, according to Murray (1990), the (La/Ce) N values of sedimentary rocks in mid-ocean ridges are ≥ 3.5, while those in ocean basins (or deep-sea plains) are 2–3, those in continental margin basins are 1–2, and those in craton basins are ≤ 1. The (La/Ce) N eigenvalue of DSDMM is 0.951–1.462 with an average of 1.209 (Table 5). According to the identification criteria of Murray et al. (1990), this shows that the DSDMM was formed in the continental margin basin.
4.3.3 (La/Yb) N ratio
This ratio is an eigenvalue describing the degree of HREE differentiation in rocks and is also an important index for judging the nature of sedimentary basins. The average (La/Yb) N value of sedimentary rock near a mid-ocean ridge is about 0.3, those of oceanic basins (or deep-sea plains) are 0.3–1.1, those of continental margin basins are 1.1–1.4, and those of craton basins are ≥ 1.4. The range of variation in (La/Yb) N eigenvalues in the DSDMM was medium to low, ranging from 0.285 to 1.189 but mainly between 0.4 and 0.7 with an average of 0.592 (Table 5). This indicates that the degree of HREE differentiation is medium to low. According to the identification standard of Murray et al. (1990), this reflects that the DSDMM formed in the transition zone between the ocean basin and continental marginal basin. Although it is slightly different from the discrimination result of the (La/Ce) N eigenvalue, it is consistent with the paleo-geographic location of the micrite mound, which is developed in a transition zone between a deep-water basin and continental slope.
4.3.4 δCe anomaly
As Ce3+ can be oxidized to Ce4+, which has higher solubility in oxidizing water, it can easily be migrated and depleted, causing a negative anomaly of δCe < 1. This is often used to judge the oxidation-reduction conditions of a depositional environment (Murray et al., 1990). The δCe value of the DSDMM varies very little, ranging from 0.80 to 1.01 and mainly between 0.85 and 0.95 with an average of 0.881 (Table 5). This reflects a weak negative anomaly, so the DSDMM deposition occurred in an oxygen-poor to oxygen-deficient reducing environment, which is consistent with the paleo-oxygen analyses of δU and the U/Th and V/(V+Ni) ratios.
4.3.5 δEu anomaly
Eu3+ is easily migrated and depleted in low-temperature reducing environments, resulting in a negative anomaly, while it is readily oxidized to insoluble Eu4+ in oxidative or high-temperature environments, resulting in relative enrichment and a positive anomaly. Therefore, δEu can be used to determine the characteristic parameters of oxidation-reduction conditions in sedimentary environments (Murray et al., 1990) and also as a basis for judging hydrothermal exhalation-sedimentation effects. The δEu value of DSDMM dolomite is very small, ranging from 0.13 to 0.32 and mainly between 0.25 and 0.30, with an average of only 0.26 (Table 5), indicative of a strong negative anomaly. Because the deposition of DSDMM occurred in a deep seafloor at low temperature, the hydrothermal fluid ejected from the seabed had a strong temperature homogenization effect with normal seawater. Therefore, this strong negative δEu anomaly indicates that the deposition of silicic dolomite occurred in an oxygen-poor to oxygen-deficient reducing environment, which is highly consistent with the analysis of paleo-oxygen facies with eigenvalues such as δCe.
4.4 Rare earth element partition pattern
REE partition patterns are an important basis for distinguishing between hot-water and non-hot-water sedimentary rocks and material sources (Fleet, 1983). In most cases, the REE partition patterns of submarine hydrothermal exhalation-sedimentation products are mostly gentle and slightly inclined to the left, and have the characteristics of relative LREE poverty and loss and HREE enrichment (Murray et al., 1990), while rocks with related material sources have similar REE partition patterns and higher correlations. The REE partition pattern for the DSDMM dolomite has a low, gentle, left-dip with relatively poor LREE and weak HREE enrichment (Fig. 6), which is completely different from the steep right-dip type of normal marine sedimentary rocks, where LREEs are enriched and HREEs are depleted. It is obvious that the low and gentle left-dip REE partition pattern of DSDMM dolomite is a significant indicator of its hydrothermal sedimentary genesis. Significantly, the partition pattern of the average values of DSDMM dolomite REEs is consistent with those of siliceous rocks and sedimentary tuffs in the same period (Fig. 7). Both are of the low, gentle, left-dip type with weak HREE enrichment, in which the correlation between dolomite and sedimentary tuff is 53% and that between dolomite and siliceous rock is 51%, indicating strong affinity with each other. This feature further confirms that the submarine exhalation of hydrothermal fluid that formed the siliceous dolomitic micrite mound mainly came from volcanic hydrothermal fluid that erupted in the same period.