Hindawi
Geofluids
Volume 2018, Article ID 2197891, 23 pages
https://doi.org/10.1155/2018/2197891
Research Article
The Fe-Zn Isotopic Characteristics and Fractionation Models:
Implications for the Genesis of the Zhaxikang Sb-Pb-Zn-Ag
Deposit in Southern Tibet
Da Wang,1,2 Youye Zheng
,1,3 Ryan Mathur
,2 and Song Wu1
1
State Key Laboratory of Geological Processes and Mineral Resources, School of Earth Sciences and Resources,
China University of Geosciences, Beijing 100083, China
2
Department of Geology, Juniata College, Huntingdon, PA 16652, USA
3
State Key Laboratory of Geological Processes and Mineral Resources and Faculty of Earth Resources, China University of Geosciences,
Wuhan 430074, China
Correspondence should be addressed to Youye Zheng; zhyouye@163.com and Ryan Mathur; MATHURR@juniata.edu
Received 30 August 2017; Revised 17 December 2017; Accepted 21 January 2018; Published 20 March 2018
Academic Editor: Bin Chen
Copyright © 2018 Da Wang et al. This is an open access article distributed under the Creative Commons Attribution License, which
permits unrestricted use, distribution, and reproduction in any medium, provided the original work is properly cited.
The genesis of the Zhaxikang Sb-Pb-Zn-Ag deposit remains controversial. Three different geological environments have been
proposed to model mineralization: a hot spring, a magmatic-hydrothermal fluid, and a sedimentary exhalative (SEDEX)
overprinted by a hot spring. Here, we present the electron probe microanalysis (EPMA) and Fe-Zn isotopic data (microsampled)
of four samples from the first pulse of mineralization that show annular textures to constrain ore genesis. The Zn/Cd ratios
from the EPMA data of sphalerite range from 296 to 399 and overlap the range of exhalative systems. The 𝛿56 Fe values of MnFe carbonate and 𝛿66 Zn values of sphalerite gradually decrease from early to late stages in three samples. A combination of the
EPMA and isotopic data shows the Fe-Zn contents also have different correlations with 𝛿66 Zn values in sphalerite from these
samples. Rayleigh distillation models this isotope and concentration data with the cause of fractionation related to vapour-liquid
partitioning and mineral precipitation. In order to verify this Rayleigh distillation model, we combine our Fe-Zn isotopic data with
those from previous studies to establish 12 Fe-Zn isotopic fractionation models. These fractionation models indicate the 𝛿56 Fe𝑖 and
𝛿66 Zn𝑖 values (initial Fe-Zn isotopic compositions) of the ore-forming system are in the range of −0.5‰∼−1‰ and −0.28‰∼0‰,
respectively. To conclude, the EPMA data, Fe-Zn isotopic characteristics, and fractionation models support the SEDEX model for
the first pulse of mineralization.
1. Introduction
To date, the Zhaxikang Sb-Pb-Zn-Ag deposit is the only
super large deposit that has been identified within the
North Himalayan Polymetallic Metallogenic Belt (NHMB).
Although basic research including the geology, petrography, geochronology, and geochemistry studies has been
conducted (e.g., [1, 2]), the genesis of this deposit is
still debated due to the complicated mineralogy and the
presence of multiple stages of mineralization. The main
viewpoints involve a hot spring [3, 4], two magmatichydrothermal fluids [5, 6], and a SEDEX overprinted by
hot spring [7] genetic models. However, most of these
genetic models are based on the S, C, O, H, and Si
isotopic evidence, which cannot absolutely trace the metal
source.
The traditional light stable C, H, O, S, and N isotopes
have been widely used to constrain fluid evolution and
metal sources in ore deposit studies (e.g., [8, 9]). However,
the evidences for metal source from these elements are
always indirect and putative as they are not the metallogenic
elements themselves [10]. For instance, these elements usually
have different characteristics with the changing of tectonic
settings, and sometimes they may even have different sources
from the metallogenic elements [11]. However, the nontraditional transition metal stable isotopes (e.g., Fe, Zn, Cu, Cd,
Mg, Cr, Sn, and Mo) are more precise tracers for the metal
sources and ore-forming processes in metallogenic systems
2
[10, 12, 13]. The development of Multicollector-Inductively
Coupled Plasma Mass Spectrometer (MC-ICP-MS) technology has greatly improved the precision of isotopic analyses
[14, 15], which results in the wide application of the nontraditional transition metal stable isotopes in economic geology
studies (e.g., [16, 17]).
The Fe-Zn isotopes are two of the most representative
isotopes applied in ore deposit studies. For example, Mason et
al. [18] and Wilkinson et al. [19] both identified that the 𝛿66 Zn
values of minerals precipitating from the same hydrothermal
fluid become heavier over time by studying the Zn isotopic
fractionation of the Alexandrinka volcanic hosted massive
sulfide (VHMS) type deposit in Russia and Midlands Irishtype deposit in Ireland, respectively. The gradual increasing
𝛿66 Zn values both from early to late stages and from south
to north within the Red Dog ore district in Alaska record
the temporal-spatial evolution of the ore-forming fluid and
constrain the SEDEX genesis with a single Zn source [20].
In addition, Fe isotopic studies of skarn Cu−Au±Fe deposits
in South China excluded the possibility that Fe originated
from sedimentary strata [21, 22]. These Fe isotopes matched
the igneous source rocks and mineralization, and the 𝛿56 Fe
values of sulfides gradually increase both from early to late
stages and away from the ore-related igneous rocks. Wang
et al. [21, 22] also revealed that the Fe isotopes fractionate
during fluid exsolution and that the ore-forming fluid is
enriched in light isotopes relative to ore-related igneous
rocks. To the contrary, Wawryk and Foden [23] investigated
the Fe-isotope fractionation in the Renison Sn-W deposit
in Australia and discovered that Fe isotopic compositions
of pyrite (0.61‰∼1.14‰), chalcopyrite (0.18‰∼0.71‰),
and magnetite (0.50‰∼0.70‰) are isotopically heavier than
Renison granite (0.27‰∼0.45‰) and thus hypothesized that
a magmatic-hydrothermal fluid exsolved from an isotopically
heavy reduced magma could deposit isotopically heavy ore
minerals whereas oxidized magmas crystallise magmatic
magnetite could result in an isotopically lighter melt and
fluid. These studies demonstrate the potential of Fe-Zn
isotopes to trace the metal source and provide insights into
ore-forming evolution.
With regard to the Zhaxikang deposit, Duan et al. [5]
have investigated the Zn isotope of sphalerite, galena, FeMn carbonates, and igneous rocks and speculated that the
𝛿66 Zn values of the hydrothermal fluid are 0.39 ± 0.10‰.
This value is consistent with those of basement rocks (average
value of 0.36‰ ± 0.03‰) and Fe-Mn carbonates (average
value of 0.27‰ ± 0.15‰), which is identified as the evidence for the magmatic origin. Meanwhile, the contribution of regional sedimentary rocks is conjectured by the
Zn-Pb-S isotopes: (1) the Zn isotopic variation range of
sulfides (−0.25‰∼0.03‰) is larger than basement rocks
(0.05‰∼0.21‰); (2) the radiogenic Pb isotopic compositions of sulfides (e.g., 206 Pb/204 Pb = 18.727∼19.896) is higher
than regional igneous rocks (206 Pb/204 Pb = 18.4∼19.2); (3) the
𝛿34 S values of sulfides (6‰∼12‰) are lighter than regional
sedimentary wall rocks (10.94‰∼11.49‰) but higher than
mantle value (0 ± 2‰). However, the Zn isotopic fractionation during the fluid exsolution and leaching process
Geofluids
[27, 28] has been ignored in Duan et al. [5]. In addition,
Wang et al. [26] also studied the Fe-Zn isotopes of the pyrite,
sphalerite, and Mn-Fe carbonate in Zhaxikang deposit, which
successfully constrained the two pulses of mineralization by
the temporally increasing 𝛿56 Fe and decreasing 𝛿66 Zn values
recorded in the deposit that coincided with an increase in
alteration. The Fe-Zn isotopic research also demonstrated the
magmatic-hydrothermal fluid origin of the second pulse of
mineralization by the heavier 𝛿56 Fe values of stage 3 pyrite
and excluded the possibility that slate is the metal source by
the similar 𝛿66 Zn values of slate and sphalerite. Nevertheless,
the attempt to trace the metal source of the first pulse of
mineralization failed in Wang et al. [26]. In this study, we
present the Fe-Zn isotopic values and variations within four
annular polished section samples from Zhaxikang deposit to
provide more credible evidence for the primary and earlier
stages of ore genesis.
2. Geological Setting
2.1. Regional Geology. The NHMB is in the eastern section of
the North Himalayan Tectonic Belt (NH), in the Himalayan
terrane. From north to south, the Himalayan terrane is
divided into four tectonic belts: the North Himalayan Tethys
Sedimentary Fold Belt (the North Himalayan Tectonic Belt;
NH), the High Himalayan Crystalline Rock Belt (HH), the
Low Himalayan Fold Belt (LH), and the Sub-Himalayan
Tectonic Belt (SH; Figure 1(a)) [29–31]. These belts are
separated by four nearly EW-trending faults, including the
South Tibet Detachment System (STDS), the Main Central
Thrust (MCT), the Main Boundary Thrust (MBT), and
the Main Frontal Thrust (MFT; Figure 1(a)) [24, 32]. The
NH, composed of a set of Palaeozoic marine sedimentary
sequences that formed in a passive continental margin
environment within northern India [33], is located between
the Indus-Yarlung Zangbo Suture Zone (IYZS) and the HH
(Figure 1(a)).
The sedimentary sequence in the NH records Late Precambrian to Devonian prerift, Carboniferous to Early Jurassic
syn-rift and Middle Jurassic to Cretaceous passive continental
margin sediments (Figure 1(b)) [66–68]. These sediments
crop out in an EW and NWW trend and are predominated
by the Precambrian Laguigangri Group and a series of
Upper Triassic, Jurassic, Lower Cretaceous, and Quaternary
sediments. The Laguigangri Group crops out in the core of
the Yelaxiangbo dome and is composed of schist, gneiss,
and migmatite units (Figure 1(b)). A set of Late TriassicEarly Cretaceous flysch formations deposit in neritic-bathyal
environments and crop out across the NH. This set of
formations are dominated by turbidite deposits and host
the majority of the Au-Sb-Pb-Zn-Ag deposits in the NH.
The lithology of these Late Triassic-Early Cretaceous flysch
formations is weak-metamorphic slate that is intercalated
with metamorphosed fine-grained sandstone, argillaceous
limestone, micrite, and siliceous rock that is intercalated
with volcanic rocks [7]. Some quaternary sediments also
occur in the central and southern area of the NH, which are
composed of gravel, sand gravel, sandy loam, clay, and ice
boulder.
Geofluids
3
90∘
91∘
92∘
93∘
N
(b)
IYZS
(km)
Lhasa Terrane
Jiacha
Ranba dome
29
Baijia
Haweng
∘
Jiangzi
Langkazi
Kangbugunba
Shalagang
Qusong
Wuladui
Qiangtang Terrane
Rongbu-Gudui Fault
Jibupu
(km)
BNSZ
Cheqiongzhubu
Lhasa Terrane
Lhasa
IYZS
STDS
LH
Cuomei
Rangla
Xiaba
30∘
Jiena
NH Fig. (b)
Keyue
Longzi
Zedang
Jiangzhula
Jisong
Langzhang
Zhaxikang
D
ST
Cuona
Bhutan
Cretaceous: clastic and carbonate rocks,
interlayered with volcanic rocks
Jurassic: shelf-facies shale (slab), sandstone,
limestone, interlayered with volcanic rocks
Upper Triassic:deep-water shelf-facies sandstone
shale (slab), interlayered with volcanic rocks
Lower-middle Triassic: shallow-marine
sandstone and shale (slab)
Permian:conglomerate, sandstone,
marble, and silty slate
Gudui
Luo
zha
faul
t
HH
MCT MBT
SH
India continent
MFT
88∘
92∘
Mazhala
91
∘
29∘
Chalapu
Zhegu
Cimalong
200
0
Xigong
Zhemulang
Lazi-Qiongd
uojiang fau
lt
Xiaqiong
Kangma dome
(a)
Bangbu
Xiangdala
Quexiaxiongqu
Yelaxiangbo dome
Kaliepu
Delong
25
0
28∘
92
∘
93
∘
Precambrian
metamorphic rocks
Late Jurassic-Early
Cretaceous mafic rocks
Thrust fault
Cenozoic granitoids
Fault
Different grade fault
Gological boundary
Pb-Zn deposit
Ductile shear zone
Internation border
Sb-Pb-Zn(Ag) deposit
Detachment fault
County town or city
Au deposit
Sb deposit
Sb-Au deposit
Figure 1: (a) Tectonic framework of the Himalayan Terrane (modified from Yin [24]). (b) Regional geological map of the North Himalayan
Polymetallic Metallogenic Belt (modified from Zheng et al. [25]). BNSZ: the Bangong-Nujiang Suture Zone; IYZS: the Indus-Yarlung Zangbo
Suture Zone; STDS: the South Tibet Detachment System; MCT: the Main Central Thrust Fault; MBT: the Main Boundary Thrust Fault; MFT:
the Main Frontal Thrust; NH: the North Himalayan Tethys Sedimentary Fold Belt; HH: the High Himalayan Crystalline Rock Belt; LH: the
Low Himalayan Fold Belt; SH: the Sub-Himalayan Tectonic Belt.
The EW-trending and NS-trending faults, both with the
several episodes of motion, are present in the NH. The EWtrending faults, controlling the distribution of intermediateacid magmatic rocks and ore deposits in the NH [7], are
older and cover a larger area than the NS-trending faults.
The representative EW-trending faults include the LaziQiongduojiang, Rongbu-Gudui, and Luozha faults as well as
the STDS and numerous metamorphic core complexes (Figure 1(b)). A series of rifts that formed from 25 Ma to present
are associated with these faults [29, 69–71], mainly including
the Sangri-Cuona, Yadong-Gulu, Shenzha-Xietongmen, and
Dangreyongcuo-Gucuo rift zones from east to west [72]. In
addition, the NS-trending faults that are considered as the
result of east-west extension of the Qinghai-Tibet Plateau
[73, 74] also formed during this period especially from 18 to
4 Ma [75, 76]. These NS-trending faults are also the important
ore-controlling structures in the NH [1].
Magmatism in the NH primarily includes the Mesozoic
and Cenozoic magmatism. The Mesozoic magmatism generated multiple suites of mafic-intermediate igneous rocks
between the Late Triassic and the Early Cretaceous, including
basaltic volcanic interlayers, dyke swarms, and subvolcanic
dykes. According to the previous geochronological data,
the SHRIMP U-Pb ages of the basic dyke swarms from
different area in the NH are 134.9 ± 1.8 Ma, 135.5 ± 2.1 Ma
[77], and 138.0 ± 3.5 Ma [78], respectively. The SHRIMP
U-Pb age of the gabbro is 155.8 Ma [79]. Tong et al. [78],
Pan et al. [80], and Zhong et al. [81] regarded these basic
dyke swarms as the result of late-stage massive expansion of
Neo-Tethys Ocean under the structural environment of the
Himalaya passive continental margin intensive stretching and
breaking-off, lithosphere extension-thinning, and asthenosphere upwelling. On the contrary, Zhu et al. [82] and Qiu et
al. [83] suggested that these basic dyke swarms are the result
of interaction between mantle plume and lithospheric mantle
material and form in the continental-rift environment. The
Cenozoic magmatism is characterized by the formation of
monzogranite, leucogranite, diorite, porphyritic diorite, and
aplite units [84, 85]. These Cenozoic intermediate-acidic
intrusive masses are distributed along the EW-trending faults
and in the core of Ranba, Kangma, and Yelaxiangbo dome
in the form of batholith, laccolith, and dykes (Figure 1(b)).
This phenomenon is considered to be the result of crustal
thickening [86] related to the collision of the India Plate and
the Eurasia Plate during the postcollision stage (25 to 0 Ma)
[87, 88].
4
The NHMB contains many Sb, Au, Sb-Au, Pb-Zn, and
Sb-Pb-Zn-Ag deposits, and the Zhaxikang Sb-Pb-Zn-Ag, the
Mazhala Au-Sb, the Chalapu Au, the Bangbu Au, the Shalagang Sb, the Cheqiongzhuobu Sb deposits are representative
(Figure 1(b)) [7, 25]. The geneses and metallogenic age of
these deposits are controversial due to the limited research,
the genetic models mainly include the SEDEX overprinted
by hot spring, carlin and carlin-like, hot spring, subvolcanic
magmatic-hydrothermal fluid, and orogenic types [25].
2.2. Ore Deposit Geology. The Zhaxikang Sb-Pb-Zn-Ag deposit is located ∼48 km west from Longzi County Town
within the southeastern Yangzuoyong-Nariyong composite
syncline in the NH (Figure 1(b)). This deposit has a reserve
of 1.268 Mt Pb-Zn with a 3.66% average Zn grade and a
2.45% average Pb grade, 0.138 Mt Sb with an average grade of
1.08%, 1800 t Ag with an average of 99.55 g/t, 3.9 t associated
Au, 361 t associated Ga, and 20 Mt Mn-Fe carbonate ores
with an average grade of 42% for Fe + Mn [89], which
makes it the largest deposit within the NHMB. The majority
of mineralization in the orefield is hosted by the Lower
Jurassic Ridang formation that consists of epimetamorphic
marine clastic rocks. This formation, dipping shallowly to
the north and striking eastwest, is divided into five lithologic
sections (Figure 2(a)). A few Upper Jurassic Weimei formations composed of fine-grained metamorphic quartzose
sandstone, silty slate, and calcarenite as well as Quaternary
sediments distributed along valleys also crop out in the
orefield (Figure 2(a)) [7].
The Zhaxikang deposit developed extensive geological
structures. A near northsouth striking fault system is prevalent in the orefield, which coexists with a group of northeaststriking faults and some folds. Engineering and geological
mapping projects have identified 16 faults, the majority of
which are steeply dipping normal faults associated with
tensional stress and wrench faults associated with torsional
stress. Faults F2, F4, F5, F6, F7, F13, F14, and F16 are the main
ore-bearing faults, faults F1 and F10 are partly mineralized,
fault F3 was associated with late-stage mineralization, faults
F8 and F9 are wrench faults without any mineralization, and
fault F12 is a nonmineralized regional fault (Figure 2(a)).
The orebodies I–VI are hosted by nearly NS-striking faults
and orebodies VII–IX are hosted by nearly NE-striking faults
(Figure 2). Our samples in this study are all from the orebody
V, which is the largest and richest one among these orebodies
within the orefield and hosts more than 80% of the reserves.
This orebody is >1400 m long, 1 to 30 m wide, and controlled
by fault F7 (Figure 2).
The magmatism in the orefield is associated with diabase,
porphyritic rhyolite, basalt, and leucogranite units as well
as some granite porphyry dykes that intruded into the
porphyritic rhyolite (Figure 2(a)). The diabase is identified
by drillholes and footrill in the central part of the orefield
as dykes that emplaced into the Jurassic Ridang Formation
and has been dated by zircon U-Pb methods to ∼133 Ma [7].
The rhyolite porphyry with the zircon SHRIMP U-Pb age of
∼135 Ma crops out in the western part of the orefield [90] and
the leucogranites crop out in the southern part of the orefield
over an area of <1 km2 . Additionally, the basalt usually occurs
Geofluids
near the orebody in the form of consequent layer or shear
layer distributed in slate and the contact region of slate and
diabase.
Various types of alteration associated with mineralization
have occurred in the orefield, including (1) the silicification
that is associated with Sb mineralization and generally
located in fault zones in the form of quartz veins, radiating
quartz, and quartz druse; (2) the carbonatization that is
associated with Pb-Zn mineralization in the form of MnFe carbonate veins and also formed the postmineralization
calcite; (3) the chlorite alteration that is generally confined
to massive and stellated aggregates of chlorite within diabase;
(4) the weak sericite alteration that is associated with chlorite
alteration and barren quartz; and (5) the clay alteration that
developed along the edges of fracture-related crushed zones.
Furthermore, the ore-forming elements display a vertical
sequence that is zoned from a lowermost Zn (Pb + Ag) zone
through a central Zn + Pb + Ag-(Sb) zone to an uppermost
Pb + Zn + Sb + Ag zone, although no horizontal zoning is
present [89].
2.3. Ore Paragenetic Sequence. The paragenetic sequence in
the Zhaxikang deposit is divided into six stages of ore formation based on the detailed hand specimen and microscopic
observations. These six stages are assigned to two clear pulses:
the first pulse consists of stages 1 to 2 and is characterized
by the assemblages of Mn-Fe carbonates and sulfides, and
the second pulse includes stages 3 to 6 and is primarily
dominated by quartz, calcite, sulfosalt minerals, and sulfides
(Figure 3).
Stage 1, dominated by a Mn-Fe carbonate + sphalerite
+ pyrite + arsenopyrite assemblage, is the initial stage of
ore formation in the Zhaxikang deposit. Majority of the
fine-grained sphalerite, pyrite, and arsenopyrite are hosted
by fine-grained Mn-Fe carbonate in the form of laminae
(Figures 4(a)–4(c)), and a few sulfides occur within the MnFe carbonates as stellated aggregates (Figure 4(d)). The finegrained layered and colloform with synsedimentary features
(Figure 4(b)). The laminae and Mn-Fe carbonates in some
samples have been cut by later stage 4 quartz-boulangerite
veins (Figure 4(c)) or have been affected by the stage 2 coarsegrained sphalerite (Figure 4(a)).
Stage 2 hosts majority of the Pb-Zn mineralization in
the Zhaxikang deposit and comprises an assemblage of MnFe carbonate + galena + sphalerite + pyrite ± arsenopyrite.
The more abundant and coarser-grained sulfides are hosted
by coarse-grained Mn-Fe carbonate and slate to form the
banded (Figure 4(k)), net-veined (Figure 4(i)), massive (Figure 4(h)), concentric annular (Figures 4(g) and 5), and Dal
Matianite (Figure 4(l)) ores. The Mn-Fe carbonate during
this stage recrystallized to different degree (Figures 4(f), 4(g)
and 4(i)–4(k)), some even formed the druse containing the
idiomorphic columnar quartz, needle-like boulangerite, or
valentinite (Figure 4(j)). We can also observe that the later
sphalerite replaces the earlier pyrite containing automorphic
stage 1 arsenopyrite to form a skeletal texture during this stage
(Figure 4(w)). Zheng et al. [7] considered that the ore textures
in stages 1 and 2 are similar to those of the Red Dog SEDEXtype ore district in Alaska.
Geofluids
5
F16
91∘ 59
92∘ 00
IX
92∘ 01
N
F7
28∘ 23
F5
ZK708
F1
F3
ZK705 ZK704
ZK701
ZK706
ZK721
ZK722
F6
ZK703
Line 7
ZK702
50
265
35
270
F4
28∘ 23
F2
62
270
58
270
37
270
61
270
III
42
300
IV
V
II
VI
28∘ 22
28∘ 22
(m)
0
200
J1 r5
Sandstone inserted
with slate
J1 r4
Slate inserted with
sandstone/basalt
J1 r3
Sandstone
J1 r2
Slate
J1 r
1
Metamorphic
quartz-sandstone
Rhyolite porphyry
Diabase vein
Quaternary
28∘ 21
Ore body
and number
II
F3
Geological
boundary
42
300
Line 7
ZK701
56
315
Fault number
>CJ ;HAF?
>CJ
55
315
Exploration line and
drill hole number
91∘ 59
VII
28∘ 21
F13
F14
VIII
92∘ 01
92∘ 00
(a)
93∘
4883
4800
ZK708 ZK706
ZK705
ZK703
ZK704
ZK701
III
4700
200
(m)
ZK702 ZK721
ZK722
VII
VI 120.02
PD7
CM710
200
V
Altitude: 4575m
4500
m
.4 9 m
IV
4600
(m)
0
307.43 m
312.16 m
323.58 m
4400
473.68 m 464.96 m
4300
4200
664.75 m
4100
4000
3950
883.7 m
ZK701
Orebody
Drill hole
PD7
CM710
Trench
Sampling location
(b)
Figure 2: (a) Geological map of the Zhaxikang Sb-Pb-Zn-Ag polymetallic deposit (modified from Zheng et al. [7]). (b) Cross-section along
Exploration Line 7. 𝐽1 𝑟1 to 𝐽1 𝑟5 : the first to fifth lithologic section of the Lower Jurassic Ridang formation.
Stage 3, characterized by the formation of a quartz ±
calcite + pyrite + sphalerite + galena ± chalcopyrite ±
arsenopyrite assemblage without Mn-Fe carbonate, is the
earliest stage of the second pulse of mineralization. The massive, veined, net-veined, and brecciated sphalerite, galena,
and pyrite occur in the quartz and calcite (Figures 4(l)–4(p),
4(x)), and most of the sulfides form by the modification
of sulfides from earlier stages. Some chalcopyrite grains
are distributed in sphalerite, galena, and pyrite as stellated
aggregates (Figures 4(y) and 4(z)). Some of these sulfides have
been cross-cut by the later quartz-boulangerite or quartzcalcite veins (Figure 4(o)).
Stage 4 is marked by a mineral assemblage composed of
quartz + antimony-lead-silver sulfosalt minerals (Figure 4(r))
that prevailingly include boulangerite and jamesonite, as well
as minor bournonite, tetrahedrite, and andorite. This stage
hosts the majority of the Sb and Ag mineralization and yields
the ores with relatively high average Ag grades. The minerals
6
Geofluids
Supergene
Stage Stage 1
Stage 2 Stage 3 Stage 4 Stage 5 Stage 6 stage
Mineral
Mn-Fe
carbonate
Sphalerite
Galena
Pyrite
Arsenopyrite
Chalcopyrite
Quartz
Calcite
Sericite
Boulangerite
Jamesonite
Bournonite
Zinckenite
Freibergite
Andorite
Tetrahedrite
Stibnite
Cinnabar
Ferrihydrite
Smithsonite
Sardinianite
Valentinite
Travertine
Malachite
Siliceous sinter
Notes:
Abundant
Intermediate
Minor
Figure 3: Mineral paragenesis within the Zhaxikang Sb-Pb-Zn-Ag
deposit (modified from Wang et al. [26]).
formed in earlier stages are replaced and cross-cut by the
quartz-boulangerite veins and boulangerite of this stage
(Figures 4(c) and 4(aa)). Some samples also contain quartz
druse filled with needle-like boulangerite (Figure 4(q)).
Stage 5 is distinguished by the formation of a quartz
+ stibnite + cinnabar assemblage and hosts part of the Sb
mineralization within the deposit. Elongate-radial stibnite
and massive stibnite-cinnabar are hosted by quartz (Figures
4(s) and 4(t)). Some stibnite cross-cut the stage 4 boulangerite
(Figure 4(aa)).
Stage 6, representing the youngest stage of mineralization
in the Zhaxikang deposit, is identified by a quartz ± calcite
assemblage without sulfides. The quartz-calcite veins of this
stage cross-cut the earlier formed minerals (Figures 4(o)
and 4(ab)). Zheng et al. [7] regarded the ore textures in the
second pulse of mineralization as typical hot spring type
metallogenic features.
Supergene stage in the Zhaxikang deposit consists of
ferrihydrite, smithsonite, sardinianite, valentinite, travertine,
malachite, and siliceous sinter (Figures 4(u) and 4(v)).
3. Sampling and Analytical Methods
3.1. Sampling. The sampling points for EPMA and Fe-Zn
isotopic analyses (the powders are sampled by microdrill)
are all in the annular polished section samples 9-3, 9-8,
ZXK-1, and ZXK-2. The specific numbers, locations, and
photomicrographs of these sampling points are given in
Figures 5 and 6, respectively. These four samples are all from
the first pulse of mineralization, only the sample ZXK-2 has
been cut by a stage 3 quartz vein (Figure 5(d)).
3.2. EPMA. Chemical compositions of sulfide, Mn-Fe carbonate, and quartz were determined on a JEOL (Japan
Electron Optics Laboratory) JXA-8100 electron microprobe
(EMP) at the Second Institute of Oceanography, State
Oceanic Administration of China. The accelerating voltage is
15 kV for Mn-Fe carbonate and quartz and 20 kV for sulfide,
the beam current is 10 nA, the beam diameter is 1 𝜇m, the
secondary electronic resolution is 6 nm with the operating
distance of 11 mm, and the repeat accuracy of the sample
stage is within 1 nm. The standards are natural minerals and
synthetic oxides as those of Sun et al. [2]. The correction
program supplied by the manufacturer is used for matrix
corrections [91, 92].
3.3. Fe-Zn Isotopic Analyses. Approximately 10–50 milligrams of sample powders was placed in 15 ml Teflon jars
and the solids were dissolved in 4 ml of heated ultrapure
aqua regia. The solutions were dried and then Fe and Zn
were purified using the BioRad MP-1 anion exchange resin
using the protocol from Maréchal et al. [14]. Yields from
the columns were tested volumetrically on the ICP-OES at
Pennsylvania State University and were all greater than 95%.
Isotope values are reported in the traditional per mil values
(‰).
The Fe isotopes were measured on the Neptune MC-ICPMS at Pennsylvania State University. The instrument setup,
sample introduction, and running conditions are discussed
in greater detail in Yesavage et al. [93]. Samples were diluted
to a 3 ppm Fe solution which produced approximately a 10 V
signal on the shoulder to the argon interference peak (56 Fe
and 40 Ar16 O). Sample intensities matched the intensity of
the bracketing standard within 10%. Mass bias was corrected
for by standard-sample-standard bracketing. In-house and
international standards were measured throughout the sessions and yielded overlapping values of SRM-3126a 𝛿56 Fe
= 0.33 ± 0.08‰, 𝑛 = 8 (accepted values 𝛿56 Fe = 0.34 ±
0.1‰ 2𝜎) [93], and HPS-WU 𝛿56 Fe = 0.62 ± 0.11‰, 𝑛 = 8
(accepted values 𝛿56 Fe = 0.60 ± 0.07‰ 2𝜎) [34]. The samples
are reported relative to the international standard IRMM014 (𝛿56 Fe (‰) = [(56 Fe/54 Fe)sample /(56 Fe/54 Fe)IRMM–014 −
1] × 1000). Reported values are an average of two different
measurements and the errors fall within the range 0.1‰ 2𝜎
of the standards.
The Zn isotopes were measured on Neptune MC-ICP-MS
at Rutgers University. Correction of mass bias for Zn using
Cu (NIST 976) was employed for these samples as suggested
in [94–97] and the corrected values were then bracketed
by the standards. The samples are reported relative to the
newly developed Zn isotope standard (AA-ETH; 𝛿66 Zn (‰)
= [(66 Zn/64 Zn)sample /(66 Zn/64 Zn)AA–ETH − 1] × 1000) and all
the quoted data from previous literatures in this paper are
converted relative to the AA-ETH standard (𝛿66 ZnAA–ETH =
𝛿66 ZnJMC 3–0749 L − 0.28‰) [98]. We also compared the new
Geofluids
7
Sp2
Sp2
Sp2
Mcar2
Py2
Mcar2
Apy1-Py1-Sp1
Apy1-Py1-Sp1
Gn2
Mcar1
Py2
Mcar1-Sp1
Sp2
Mcar1
Gn2
Apy1-Py1-Sp1
Qtz4-Blr4 Vein
2 =G
4 =G
(a)
3 =G
3 =G
(b)
(c)
(d)
Py2
Mcar2
Mcar2
Py2
Sp2
Gn2
Py2
Sp2
Slate
4 =G
2 =G
3 =G
(f)
(e)
5 =G
(h)
(g)
Mcar2
Py2
Sp2
Sp2
Py2
Py2
Mcar2
Mcar2
Sp2
Mcar2
Sp2
Sp2
Qtz5-valentinite
Qtz3
4 =G
3 =G
2 =G
(j)
(i)
(k)
(l)
Sp3
Slate
Sp3
Cal3
Qtz6-Cal6
Qtz3-Cal3 Sp3
Sp3-Py3
Qtz4-Blr4
Gn3
Qtz3
2 =G
3 =G
(m)
2 =G
(o)
(n)
Blr4
(p)
Qtz5
Qtz4
Stb5
Stb5
Qtz4
Blr4
2 =G
Blr4
Ci5
5 =G
3 =G
(q)
3 =G
(r)
(s)
Qtz5
3 =G
(t)
Sp3
siliceous sinter
sardinianite
Lm
Sp2
Qtz3
Py2
Apy1
2 =G
(u)
5 =G
(v)
100 G
100 G
(w)
Figure 4: Continued.
Mcar2
(x)
8
Geofluids
Blr4
Stb5
Sp3
Stb5
Ccp3
Ccp3
Qtz6
Sp3
Py3
Qtz5
20 G
(y)
100 G
200 G
(z)
(aa)
200 G
(ab)
Figure 4: Hand specimen photographs and photomicrographs of representative samples from the Zhaxikang deposit. (a) Stage 1 lamellar
sphalerite-pyrite-arsenopyrite and stage 2 massive sphalerite-pyrite hosted within fine-grained Mn-Fe carbonate. (b) Stage 1 lamellar and
stage 2 banded Mn-Fe carbonate-sphalerite-galena ore with visible synsedimentary features including rhythmic sedimentation in the upper
part and angular folding in the lower part of the sample. (c) Stage 1 lamellar sphalerite-pyrite-arsenopyrite and stage 2 massive and banded
sphalerite-pyrite hosted by fine-grained Mn-Fe carbonate. The mineral assemblage is in turn cross-cut by stage 4 quartz-boulangerite veins.
(d) Coarse-grained stage 2 Mn-Fe carbonate-sphalerite formed by the recrystallization of fine-grained stage 1 Mn-Fe carbonate-sphalerite. (e)
Stage 2 massive coarse-grained pyrite hosted by slate. (f) Massive and brecciated stage 2 sphalerite hosted in stage 2 Mn-Fe carbonate. (g) Stage
2 massive, globular, and concentric annular sphalerite-pyrite hosted by coarse-grained Mn-Fe carbonate. (h) Stage 2 massive galena and pyrite.
(i) Stage 2 massive and veined sphalerite-pyrite hosted by coarse-grained Mn-Fe carbonate. (j) Stage 1 lamellar sphalerite-pyrite-arsenopyrite
and stage 2 massive sphalerite hosted by fine-grained Mn-Fe carbonate. The sample also contains a Mn-Fe carbonate druse dominated by
idiomorphic columnar quartz and valentinite. (k) Stage 2 coarse-grained sphalerite-pyrite hosted by stage 2 Mn-Fe carbonate with banded
textures. (l) Stage 2 sphalerite and Mn-Fe carbonate ore with typical Dal Matianite texture. (m) Disseminated stage 3 sphalerite and pyrite
hosted in stage 3 quartz and calcite, and stage 3 quartz and calcite cut the slate. (n) Stage 3 brecciated sphalerite within stage 3 quartz-calcite.
(o) Stage 3 sphalerite-galena veins cross-cut by stage 6 quartz-calcite veins. (p) Stage 3 sphalerite cross-cut by stage 4 quartz-boulangerite
veins. (q) Stage 4 massive and needle-like boulangerite hosted by stage 4 quartz. (r) Stage 4 boulangerite-quartz. (s) Stage 5 elongate stibnite
hosted by stage 5 quartz. (t) Stage 5 stibnite-cinnabar hosted by stage 5 quartz. (u) Siliceous sinter formed during the Supergene stage. (v)
Ferrihydrite and sardinianite formed during the Supergene stage. (w) Stage 2 pyrite containing automorphic stage 1 arsenopyrite is replaced
by later stage 2 sphalerite to form a skeletal texture. (x) Stage 3 sphalerite occurs in stage 3 quartz. (y) The emulsion-like and disseminated
stage 3 chalcopyrite grains are dotted in the stage 3 sphalerite. (z) The stage 3 chalcopyrite grains are dotted among the grains of stage 3 pyrite.
(aa) Stage 3 sphalerite replaced by stage 4 boulangerite that is in turn cross-cut by stage 5 stibnite. (ab) Stage 5 stibnite cross-cut by stage 6
quartz. Mcar1 = stage 1 fine-grained Mn-Fe carbonate; Apy1 = stage 1 lamellar arsenopyrite; Py1 = stage 1 lamellar pyrite; Sp1 = stage 1 lamellar
sphalerite; Mcar2 = stage 2 coarse-grained Mn-Fe carbonate; Apy2 = stage 2 arsenopyrite; Py2 = stage 2 pyrite; Sp2 = stage 2 sphalerite; Gn2
= stage 2 coarse-grained galena; Py3 = stage 3 pyrite; Sp3 = stage 3 sphalerite; Gn3 = stage 3 galena; Ccp3 = stage 3 chalcopyrite; Qtz3 = stage
3 quartz; Cal3 = stage 3 calcite; Blr4 = stage 4 boulangerite; Qtz4 = stage 4 quartz; Stb5 = stage 5 stibnite; Ci5 = stage 5 cinnabar; Qtz5 = stage
5 quartz; Cal6 = stage 6 calcite; Qtz6 = stage 6 quartz; Lm = Supergene stage ferrihydrite.
standard relative to IRMM 3702 and obtained a 𝛿66 Zn =
0.03‰, which is within the error reported in Archer et al.
[98]. Solutions were kept at 100 ppb Cu and 150 ppb Zn which
generated 63 Cu = 7 V and 66 Zn = 4 V. One block of 30 ratios
is reported and the average error for the standard compared
to itself throughout the session is 0.05‰ 2𝜎.
4. Results
4.1. EPMA. All the EPMA data are given in Tables 1 and 2. The
Mn-Fe carbonate contains 23.562∼31.806 wt% Fe and 27.514∼
32.232 wt% Mn, with a negative correlation between Fe and
Mn contents (Figure 7(a)), which indicates that the Mn-Fe
carbonates form by the isomorphic substitution of Fe2+ and
Mn2+ ions and have a molecular formula of (Mn0.5 Fe0.5 )CO3 .
The Fe contents are around 36 wt% for the arsenopyrite
samples and 46 wt% for the pyrite samples. The sphalerite
samples have 56.941∼60.552 wt% Zn and 5.375∼9.424 wt%
Fe with a negative correlation between these two elements
(Figure 7(b)).
4.2. Fe-Zn Isotopes. All the Fe-Zn isotopic data are given
in Table 3. The annular polished section samples have
𝛿56 FeIRMM-014 of −1.95‰∼0.43‰, with an average of
−0.50‰ ± 1.09‰ (2SD, 𝑛 = 19), and 𝛿66 ZnAA-ETH of
−0.38‰∼0.07‰ with an average of −0.25‰ ± 0.19‰ (2SD,
𝑛 = 12). The Mn-Fe carbonate and pyrite show the 𝛿56 Fe
values range from −1.95‰ to −0.59‰ (average value of
−0.97‰ ± 0.86‰; 2SD, 𝑛 = 8) and from −0.26‰ to 0.23‰
(average value of −0.07‰ ± 0.35‰; 2SD, 𝑛 = 9), respectively.
The two arsenopyrite samples exhibit the 𝛿56 Fe values of
−1.01‰ (ZXK-1-3; Figure 5(c)) and −0.18‰ (ZXK-2-2;
Figure 5(d)).
In sample 9-3, the 9-3-1, 9-3-2 (Figure 6(a)), and 9-3-3
sphalerite have the 𝛿66 Zn values of −0.12‰, −0.23‰, and
−0.31‰ with a gradually decreasing trend, the 9-3-4 and 9-35 Mn-Fe carbonate show the same trend with the 𝛿56 Fe values
of −0.59‰ and −1.95‰, and the 𝛿56 Fe value of 9-3-7 pyrite
is −0.26‰ (Figure 5(b)). Similarly, in sample 9-8, from 98-2 (−0.09‰; Figure 6(c)) through 9-8-5 (−0.23‰) to 9-83 (−0.35‰; Figure 6(d)), the 𝛿66 Zn values of sphalerite also
present a gradually decreasing trend; meanwhile the 𝛿56 Fe
values also decrease from −1.06‰ (9-8-8; Figure 6(b)) to
−1.38‰ (9-8-9) for Mn-Fe carbonate and from 0.23‰ (9-84) to 0.09‰ (9-8-7) for pyrite, although the 9-8-1 sphalerite
and 9-8-6 pyrite have the 𝛿66 Zn value of −0.17‰ and 𝛿56 Fe
values of −0.26‰ (Figure 5(a)). In sample ZXK-1, the 𝛿66 Zn
Geofluids
Table 1: EPMA data for Mn-Fe carbonate and quartz from the Zhaxikang deposit.
Sample number SiO2 (%)
9-3-4
0.015
9-3-5
0.024
9-8-8
0.002
9-8-9
–
ZXK-1-6
0.001
ZXK-1-7
0.028
ZXK-2-7
0.005
ZXK-2-8
–
ZXK-2-9
0.009
ZXK-2-10
0.015
ZXK-2-11
97.000
Na2 O (%)
0.005
0.109
0.031
0.025
0.028
0.011
0.039
0.051
0.049
–
0.026
K2 O (%)
0.007
0.027
0.007
–
0.001
0.019
–
–
0.008
–
0.025
TiO2 (%)
0.011
–
–
–
0.022
0.036
0.055
0.012
–
0.023
–
Al2 O3 (%)
0.003
0.010
–
0.008
–
0.014
–
0.010
0.037
0.012
0.219
MgO (%)
1.231
1.080
1.276
1.189
1.207
1.995
1.236
1.140
1.552
1.119
0.028
CaO (%)
2.061
2.151
1.978
1.144
5.178
2.940
5.542
4.298
1.946
4.849
0.018
FeO (%)
28.598
28.372
27.341
31.806
24.070
29.639
27.235
24.287
26.739
23.562
0.020
MnO (%)
29.840
30.582
31.338
27.514
30.717
27.943
27.604
32.694
30.968
32.232
—
SrO (%)
0.007
–
0.010
–
–
–
–
–
0.007
–
–
BaO (%)
0.057
0.002
0.022
–
–
0.017
–
0.008
0.013
0.003
–
CO2 (%)
38.164
37.643
37.993
38.316
38.775
37.358
38.285
37.500
38.673
38.185
2.664
Total (%)
99.999
100.000
99.998
100.002
99.999
100.000
100.001
100.000
100.001
100.000
100.000
Mineral
Mcar
Mcar
Mcar
Mcar
Mcar
Mcar
Mcar
Mcar
Mcar
Mcar
Qtz
Notes. (1) Mcar = Mn-Fe carbonate; Qtz = quartz; (2) for the sampling points, please see Figure 5.
9
10
Table 2: EPMA data for sulfides from the Zhaxikang deposit.
Sample number As (%)
9-3-1
–
9-3-2
–
9-3-3
–
9-3-6
0.784
9-3-7
2.244
9-8-1
0.002
9-8-2
–
9-8-3
–
9-8-5
0.009
9-8-6
0.006
9-8-7
0.015
ZXK-1-1
–
ZXK-1-2
–
ZXK-1-3(1)
39.757
ZXK-1-3(2)
2.435
ZXK-1-3(3)
–
ZXK-1-4
1.942
ZXK-1-5
2.056
ZXK-2-1
–
ZXK-2-2
39.643
ZXK-2-3
–
ZXK-2-4
39.444
ZXK-2-5
2.417
ZXK-2-6
0.031
S (%)
33.165
33.423
33.262
52.884
51.934
33.740
33.649
33.500
34.064
53.867
53.359
33.502
33.428
24.131
51.666
33.337
51.982
52.248
33.684
23.789
33.868
24.229
51.755
53.630
Se (%)
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
0.002
0.001
–
–
–
–
–
–
–
Cu (%)
0.133
0.015
–
0.005
0.002
0.009
0.011
0.006
0.089
0.033
0.471
0.012
0.036
0.006
0.018
0.014
0.008
0.008
0.061
0.002
0.025
–
–
0.107
Co (%)
0.009
0.012
0.015
0.044
0.049
0.010
0.005
0.003
0.004
0.050
0.043
0.009
–
0.043
0.067
0.012
0.046
0.054
0.010
0.050
0.012
0.041
0.045
0.047
Zn (%)
59.095
58.224
57.114
–
–
57.705
59.584
56.941
57.475
0.058
0.044
57.497
60.190
0.009
0.185
58.623
0.015
–
59.726
0.008
58.034
–
0.027
–
Sb (%)
–
–
–
–
–
–
–
–
–
–
–
–
0.002
–
–
–
–
–
–
0.196
–
0.173
–
–
Ag (%)
–
0.002
–
–
–
0.004
0.005
–
–
0.005
–
–
0.002
–
0.007
0.015
–
–
–
0.006
0.001
–
–
–
Au (%)
–
–
–
0.016
0.004
–
–
–
–
0.012
–
–
–
0.027
0.013
–
0.003
0.013
–
–
–
0.026
–
0.009
Cd (%)
0.171
0.151
0.159
–
–
0.160
0.150
0.160
0.151
–
–
0.172
0.201
–
–
0.147
–
–
0.170
–
0.196
–
–
–
Bi (%)
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
Pb (%)
–
–
–
0.134
0.132
0.018
–
–
–
0.300
0.174
–
–
0.063
0.113
–
0.196
0.137
–
0.081
–
0.061
0.131
0.168
NiO (%)
–
–
–
–
–
0.008
–
0.002
–
0.002
–
–
–
0.002
0.004
0.003
–
–
0.005
0.098
0.002
–
–
–
FeO (%)
6.985
8.060
9.405
46.585
46.044
8.827
7.112
9.424
8.901
46.539
46.518
9.142
5.810
36.340
45.877
7.747
46.004
46.163
6.904
36.245
8.714
36.424
46.227
46.654
Total (%)
99.558
99.887
99.955
100.452
100.409
100.483
100.516
100.036
100.693
100.872
100.624
100.334
99.669
100.378
100.385
99.900
100.197
100.679
100.560
100.118
100.852
100.398
100.602
100.646
Mineral
Sp
Sp
Sp
Py
Py
Sp
Sp
Sp
Sp
Py
Py
Sp
Sp
Apy
Py
Sp
Py
Py
Sp
Apy
Sp
Apy
Py
Py
Notes. (1) Py = pyrite, Sp = sphalerite, and Apy = arsenopyrite; (2) for the sampling points, please see Figure 5.
Geofluids
Geofluids
11
56 Fe values –0.26%₀
0.23‰ 0.09‰
Pyrite
66
9-8
Zn
values
Gn
3
–0.35%₀
5 Sp Sp
–0.23%₀
4
2
–0.09%₀
1
–0.17%₀
6 Slate
Sphalerite
Fe
contents
9.424%
8.901%
Zn
contents
56.941%
57.475%
7.112%
8.827%
59.584%
57.705%
8
Sphalerite
66 Zn
values
−0.31%₀
Zn
Fe
contents contents
57.114% 9.405%
2
−0.23%₀
58.224% 8.060%
1
−0.12%₀
59.095% 6.985%
9-3
7
56 Fe values
Mn-Fe Carbonate
–0.78%₀
–0.92%₀
9
Mcar
56 Fe values
–0.26%₀
Pyrite
Py 7
3
Mcar
56
Fe values
Mn-Fe Carbonate
−1.95%₀
6
5
Py
Sp
4
−0.59%₀
3 cm
3 cm
(a)
(b)
−0.17%₀
−0.07%₀
56 Fe values
Pyrite
56
Fe values
−0.05%₀ −0.22%₀ 0.12‰ Pyrite
ZXK-2
ZXK-1
Sp
3
11
9
6
56
56
Fe values
Arsenopyrite
−1.01%₀
5
Py-Sp-Apy
3
2
Sp
Fe values
Mn-Fe Carbonate
−0.78%₀
7
1
4
−0.66%₀
6
Mcar
Py
Mcar
56 Fe values
Arsenopyrite
−0.18%₀
2
5
Apy
1
Py
8
56 Fe values
Mn-Fe Carbonate
−1.04%₀
−1.06%₀
7
3 cm
3 cm
(c)
4
Sp
10
Gn Py
−0.15%₀ −0.18%₀ −0.38%₀ 66 Zn values
60.190% 58.623% 57.497% Zn contents Sphalerite
5.810% 7.747% 9.142% Fe contents
Qtz
−0.28%₀
58.034%
8.714%
−0.32%₀ 66 Zn values
59.726% Zn contents
Fe contents
6.904%
Sphalerite
(d)
Figure 5: The sampling points and Fe-Zn isotopic-elemental variations in the annular polished section samples: (a) 9-8; (b) 9-3; (c) ZXK-1;
(d) ZXK-2. Mcar = Mn-Fe carbonate; Apy = arsenopyrite; Py = pyrite; Sp = sphalerite; Qtz = quartz.
values of ZXK-1-3 (−0.18‰; Figures 6(f) and 6(m)) and ZXK1-2 (−0.15‰; Figure 6(g)) sphalerite are almost the same
and heavier than ZXK-1-1 (−0.38‰) sphalerite; the 𝛿56 Fe
values for pyrite increase from ZXK-1-3 (−0.22‰) to ZXK1-4 (0.12‰; Figure 6(e)) and ZXK-1-5 (−0.05‰); and, as for
Mn-Fe carbonate, the ZXK-1-6 (−0.66‰) has heavier 𝛿56 Fe
values than ZXK-1-7 (−0.78‰; Figure 5(c)). By contrast, in
sample ZXK-2, the ZXK-2-1 (−0.32‰; Figure 6(h)) and ZXK2-3 (−0.28‰; Figures 6(k) and 6(p)) sphalerite have almost
the same 𝛿66 Zn values; the ZXK-2-7 (−1.06‰) and ZXK-28 (−1.04‰) Mn-Fe carbonate also yield approximate 𝛿56 Fe
values; also the ZXK-2-5 (Figures 6(i) and 6(n)) and ZXK-26 pyrite show the similar 𝛿56 Fe values of −0.07‰ and −0.17‰
(Figure 5(d)).
5. Discussion
5.1. The Fe-Zn Isotopic and Elemental Variations. Sample 93 with typical concentric annular texture has the gradually
decreasing 𝛿66 Zn values of sphalerite and 𝛿56 Fe values of MnFe carbonate from core to edge (early to late; Figure 5(b)).
Similarly, sample 9-8 shows the same isotopic variation trend
of pyrite, sphalerite, and Mn-Fe carbonate except the sulfides
in the core (Figure 5(a)). As the core in 9-8 consists of
crushed, earlier formed slate breccias and sulfides, the annular sulfides and Mn-Fe carbonate formed around the core,
and thus this core is not included in the decreasing trend.
In sample ZXK-1, as the laminae (ZXK-1-3; Figure 5(c)) is
regarded as the earliest, the Fe-Zn isotopic values of sphalerite
and Mn-Fe carbonate still present gradually decreasing trend
from ZXK-1-3 (−0.18‰) to ZXK-1-1 (−0.38‰) and ZXK-1-6
(−0.66‰) to ZXK-1-7 (−0.78‰), respectively (Figure 5(c)),
whereas the 𝛿56 Fe values of pyrite show the different variation, which might be related to the influence of Fe isotopic
fractionation within pyrite-arsenopyrite-sphalerite mineral
pair (Figures 6(f) and 6(m)). However, there is no similar
decreasing Fe-Zn isotopic variation trend in sample ZXK-2,
and the modification by the stage 3 quartz vein must be the
main cause (Figures 5(d), 6(h)–6(l) and 6(n)–6(p)).
Previous studies [12, 62, 99] have proposed a Rayleigh
distillation model to explain an increasing trend in 𝛿66 Zn
values within precipitates over time for the hydrothermal
fluid. This Rayleigh distillation model is as follows: the oreforming materials derived from a single source would be
subjected to kinetic Rayleigh fractionation that would lead
to the early formed mineral precipitants being preferentially
enriched in light isotopes, as well as residual fluids and
later precipitants with heavier isotopic values, causing an
increasing trend in isotopic values within precipitants over
time. Several previous studies have used this distillation
model to explain the Zn isotopic variation within different
types of deposits (e.g., VHMS: [18]; Irish-type: [19, 100];
SEDEX: [20, 53]). Likewise, this Rayleigh distillation model
is also applicable to the Fe isotopic variation in skarn
deposits [21, 22]. However, the 𝛿66 Zn and 𝛿56 Fe values
gradually decrease from early to late stages within Zhaxikang
deposit, which cannot be explained by this distillation
model. Another Rayleigh distillation mechanism models this
12
Geofluids
Mcar
Mcar
Mcar
Slate
Sp
Sp
Sp
Py
Sp
Sp
500 G
500 G
(a)
200 G
(b)
200 G
(c)
(d)
Qtz
Mcar
Mcar
Mcar
Sp
Sp
Sp
Sp-Py-Apy
Py
500 G
500 G
(e)
500 G
500 G
(f)
(g)
(h)
Mcar
Mcar
Mcar
Qtz
Qtz
Sp
Qtz
Qtz
Apy
Py
Apy
Apy
500 G
500 G
Mcar
(i)
500 G
500 G
(j)
(k)
(l)
Qtz
Sp
Apy
Qtz
Py
Mcar
Sp
Apy
Py
Py
Apy
Mcar
Gn
Mcar
Mcar
(m)
(n)
(o)
Qtz
Sp
(p)
Figure 6: The photomicrographs and electron probe micrographs of sampling area in the annular polished section samples for EPMA and
Fe-Zn isotopic analyses. (a) 9-3-2; (b) 9-8-8; (c) 9-8-2; (d) 9-8-3 and 9-8-5; (e) ZXK-1-4; (f) ZXK-1-3; (g) ZXK-1-2; (h) ZXK-2-1; (i) ZXK-2-5;
(j) ZXK-2-2; (k) ZXK-2-3; (l) ZXK-2-4; (m) ZXK-1-3; (n) ZXK-2-5; (o) ZXK-2-2; (p) ZXK-2-3. Abbreviations are as in Figure 5.
decreasing trend: the metallogenic elements are transported
by the ore-forming system consisting of vapour and liquid
phases, and there is partitioning between vapour-liquid
phases and the ratios change with the temperature decreasing.
Then the minerals precipitate from the liquid phase of the
ore-forming system. During this period, the vapour-liquid
partitioning and mineral precipitation cause the Rayleigh
fractionation, and this Rayleigh fractionation leads to the
mineral precipitation being preferentially enriched in heavy
isotopes relative to the ore-forming system. Thus, the isotopic values of subsequent minerals are lighter and lighter
[101, 102].
This Rayleigh distillation model is supported by the
following evidence: (1) The vapour-liquid partitioning and
related isotopic fractionation for transition metal elements
(e.g., Cu and Mo) have been confirmed by previous research
in the Dahutang W-Cu-Mo ore field [103]; (2) minerals typically precipitate from the liquid phase; however, according
to the previous literature [58], in unique cases the vapour
phase containing metal can even directly condensate to form
solid phases from high-temperature ore-forming system (e.g.,
VMS and volcano related system), which can demonstrate
the existence of the vapour phase for metal elements. As
for the Zhaxikang deposit, the theoretical calculated oreforming temperature from Fe isotopic data is 500∼800∘ C
[26], and thus there should be a transitory high-temperature
period, making the vapour-liquid partitioning possible; (3)
the fluid inclusion data demonstrate that Mn-Fe carbonates
and sulfides exist in three types of inclusions: A gas-liquid
two-phase water inclusions (W type, more than 90%), B pure
13
34
61
32
60
Zn content (%)
Mn content (%)
Geofluids
30
y = −0.535x + 44.583
28
R2 = 0.5242
y = −0.909x + 65.701
R2 = 0.9624
58
57
26
24
59
56
26
28
30
Fe content (%)
32
34
5
6
7
(a)
10
(b)
0.05
0.05
0.00
0.00
−0.05
−0.10
66 Zn (% ₀)
66 Zn (% ₀)
9
Sphalerite
Mn-Fe carbonate
−0.15
−0.20
y = −0.0633x + 0.2947
−0.25
R2 = 0.6149
−0.30
−0.35
−0.40
5.00
8
Fe content (%)
6.00
7.00
8.00
Fe content (%)
9.00
10.00
Sphalerite
−0.05
−0.10
−0.15
−0.20
−0.25
−0.30
−0.35
−0.40
56.00
y = 0.072x − 4.4143
R2 = 0.6496
57.00
58.00
59.00
Zn content (%)
60.00
61.00
Sphalerite
(c)
(d)
Figure 7: (a) The negative correlation between Mn and Fe contents within Mn-Fe carbonate. (b) The negative correlation between Zn and
Fe contents within sphalerite. (c) The negative correlation between 𝛿66 Zn values and Fe contents within sphalerite from 9-3, 9-8, and ZXK-1.
(d) The positive correlation between 𝛿66 Zn values and Zn contents within sphalerite from 9-3, 9-8, and ZXK-1.
liquid inclusions (L type), and C pure CO2 inclusions (PG
type) [104] from Zhaxikang deposit.
On the other hand, for sphalerite, there are positive correlations between Zn contents and 𝛿66 Zn values, with negative
correlations between Fe contents and 𝛿66 Zn values in samples
9-8 (Figure 5(a)), 9-3 (Figure 5(b)), and ZXK-1 (Figure 5(c)),
respectively. Moreover, plotting all the data from these three
samples on a diagram together, the correlations are also
good with 𝑅2 > 0.6 (Figures 7(c) and 7(d)). Zinc and iron
usually show similar geochemical behaviour as both of them
are highly mobile in chloride-bearing hydrothermal fluids
[23, 100]. Thus, the Zn2+ and Fe2+ ions are preferentially
enriched in the liquid phase relative to the vapour phase
before precipitation [23, 100], which cause the decreasing Zn
contents of sphalerite over time. As the total content of Zn2+
and Fe2+ is constant in sphalerite, the Fe contents of sphalerite
gradually increase with the decreasing Zn contents. These
correlations further support the hypothesis that the oreforming system is the mixture of vapour and liquid phases. In
addition, the sample ZXK-2 cut by a later stage 3 quartz vein
does not present the same correlations and variations, which
is a new evidence for two pulses of mineralization proposed
by Zheng et al. [7] and Wang et al. [26].
All of the evidence above reveals that the ore-forming
elements are transported by the ore-forming system that
consists of vapour and liquid phases. The vapour-liquid partitioning and mineral precipitation are the main cause of Fe-Zn
isotopic and elemental variations. Afterwards, the overprint
by the second pulse of mineralization has also partly modified
the Fe-Zn isotopic and elemental compositions of some
earlier samples (Figure 5(d)) [26].
5.2. The Fe-Zn Isotopic Fractionation Models for the OreForming System. In order to verify the Rayleigh distillation
model in Section 5.1 and obtain more information of the oreforming system, we use the following equations to establish
Fe-Zn isotopic fractionation models for the ore-forming
system and mineral precipitation:
𝛿Minerals (‰) =
𝛼
× (𝛿𝑖 + 1000) × 𝐹(𝛼𝑚 −1)
𝛼𝑖
− 1000,
1
× (𝛿𝑖 + 1000) × 𝐹(𝛼𝑚 −1)
𝛿Ore-forming System (‰) =
𝛼𝑖
− 1000.
(1)
14
Geofluids
Table 3: Fe-Zn isotopic data for annular polished section samples from the Zhaxikang deposit.
Sample number
9-3-1
9-3-2
9-3-3
9-3-4
9-3-5
9-3-7
9-8-1
9-8-2
9-8-3
9-8-4
9-8-5
9-8-6
9-8-7
9-8-8
9-8-9
ZXK-1-1
ZXK-1-2
ZXK-1-3(1)
ZXK-1-3(2)
ZXK-1-3(3)
ZXK-1-4
ZXK-1-5
ZXK-1-6
ZXK-1-7
ZXK-2-1
ZXK-2-2
ZXK-2-3
ZXK-2-5
ZXK-2-6
ZXK-2-7
ZXK-2-8
Mineral
Sp
Sp
Sp
Mcar
Mcar
Py
Sp
Sp
Sp
Py
Sp
Py
Py
Mcar
Mcar
Sp
Sp
Apy
Py
Sp
Py
Py
Mcar
Mcar
Sp
Apy
Sp
Py
Py
Mcar
Mcar
𝛿56 FeIRMM–014
2𝜎
−0.59
−1.95
−0.26
0.08
0.08
0.08
0.23
0.08
−0.26
0.09
−0.78
−0.92
0.08
0.08
0.08
0.08
−1.01
−0.22
0.08
0.08
0.12
−0.05
−0.66
−0.78
0.08
0.08
0.09
0.08
−0.18
0.08
−0.07
−0.17
−1.06
−1.04
𝛿66 ZnAA–ETH
−0.12
−0.23
−0.31
2𝜎
0.05
0.05
0.05
−0.17
−0.09
−0.35
0.05
0.05
0.05
−0.23
0.05
−0.38
−0.15
0.05
0.05
−0.18
0.05
−0.32
0.05
−0.28
0.05
0.08
0.08
0.08
0.08
Notes. (1) Abbreviations and sampling points are as in Tables 1 and 2; (2) 2𝜎 is two times the standard deviation.
𝛿Minerals , 𝛿Ore-forming System , and 𝛿𝑖 are the 𝛿56 Fe-𝛿66 Zn values
of momentary mineral precipitation and momentary and
initial ore-forming system, respectively; 𝛼, 𝛼𝑖 , and 𝛼𝑚 are
the isotopic fractionation factors between ore-forming system and mineral precipitation that refer to the momentary
condensation temperature 𝑇, the initial temperature of oreforming system 𝑇𝑖 , and (𝑇 + 𝑇𝑖 )/2, respectively; and 𝐹 is
the fraction of remaining ore-forming system that consists
of vapour and liquid phases [102, 105]. In addition, the
equations for fractionation factors are approximately using
ln 𝛼Fe = 0.4432 × 106 /𝑇2 for Fe [106] and ln 𝛼Zn = 0.2853 ×
106 /𝑇2 + 0.0535 for Zn (𝑇 is absolute temperature in K)
[107].
Wang et al. [26] have calculated the ore-forming temperature (500∼800∘ C) of the first pulse of mineralization
in Zhaxikang deposit using the Fe isotopic fractionation
between pyrite and Mn-Fe carbonate. Although this temperature range is a little wide, the highest temperature of
such ore-forming system can reach around 500∘ C according
to the previous ore deposit studies [6, 22, 23]; hence we
regard 500∘ C as the initial temperature of the ore-forming
system. The homogenization temperature (240∘ C) of the fluid
inclusions [104] from the first pulse of mineralization in
Zhaxikang deposit is considered as the momentary condensation temperature. Furthermore, for the purpose of making
the fractionation models more comprehensive and exact, we
also quote the Fe-Zn isotopic data of pyrite (𝛿56 Fe: stage
1: −0.33‰ to −0.09‰; stage 2: −0.30‰ to 0.19‰; stage 3:
0.16‰ to 0.43‰), sphalerite (𝛿66 Zn: −0.31‰ to 0.07‰),
and Mn-Fe carbonate (𝛿56 Fe: −0.80‰ to −0.55‰; 𝛿66 Zn:
−0.11‰ to 0.04‰) from Wang et al. [26], as well as the 𝛿66 Zn
values of sphalerite (−0.25‰ to 0.03‰) and Mn-Fe carbonate
(−0.01‰) from Duan et al. [5]. Finally, we set up 12 FeZn isotopic fractionation models for pyrite, sphalerite, and
Mn-Fe carbonate (Figure 8) with different 𝛿56 Fe𝑖 values of
0‰ (mean value of magma) [51], −0.5‰, and −1‰, as well
as 𝛿66 Zn𝑖 values of −0.28‰ (the lightest value of seafloor
hydrothermal fluids) [62], 0‰ (mean value of bulk earth)
48.8%
Ore-forming system
F
1.00
Pyrite
Ore-forming system
−1.00
−2.00
48.7%
26.7%
0.07‰
−0.38%₀
Sphalerite
Ore-forming system
−3.00
−4.00
2.00
1.00
0.00
−1.00
−2.00
−3.00
−4.00
66 :Hi = 0.23%₀
Ore-forming system
F
F
−1.95%₀
7.1%
Ore-forming system
10.9%
56 &?i = −0.5%₀
Mn-Fe carbonate
−0.55%₀
39.7%
−1.95%₀
Ore-forming system
F
1.00
0.00
−1.00
−2.00
−3.00
−4.00
−5.00
−6.00
−7.00
56 &?i = −1%₀
63.3%
Mn-Fe carbonate
17.3%
Ore-forming system
F
F
(i)
(h)
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
Ore-forming system
−1.00
66 :Hi = 0%₀
46.8%
38.2%
Mn-Fe carbonate
0.04‰
−0.11%₀
−2.00
−3.00
Ore-forming system
−4.00
F
(j)
2.00
1.00
0.00
−1.00
−2.00
−3.00
−4.00
66 :Hi = 0.23%₀
34.8%
Ore-forming system
F
(k)
Mn-Fe carbonate 0.04‰
−0.11%₀
28.5%
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
Mn-Fe carbonate
0.00
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
55.6%
1.00
0.04‰
−0.11%₀
66 Zn (%₀ )
66 :Hi = −0.28%₀ 68.1%
−0.55%₀
−1.95%₀
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
−0.55%₀
1.00
0.00
−1.00
−2.00
−3.00
−4.00
−5.00
−6.00
−7.00
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
24.9%
56 Fe (%₀ )
Mn-Fe carbonate
(f)
56 Fe (%₀ )
(e)
56 &?i = 0%₀
0.07‰
−0.38%₀
36.3%
19.8%
Sphalerite
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
Ore-forming system
66 :Hi = 0%₀
0.00
0.43‰
−0.33%₀
F
66 Zn (%₀ )
Sphalerite
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
56 Fe (%₀ )
77.7%
(c)
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
−0.38%₀
38.7%
66 Zn (%₀ )
0.07‰
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
66 Zn (%₀ )
66 :Hi = −0.28%₀
70.9%
(g)
66 Zn (%₀ )
56 &?i = −1%₀
F
F
0.50
0.00
−0.50
−1.00
−1.50
−2.00
−2.50
−3.00
1.00
0.00
−1.00
−2.00
−3.00
−4.00
−5.00
−6.00
−7.00
(b)
(d)
2.00
1.00
0.00
−1.00
−2.00
−3.00
−4.00
−5.00
−6.00
0.43‰
−0.33%₀
Pyrite
(a)
0.50
0.00
−0.50
−1.00
−1.50
−2.00
−2.50
−3.00
99.0%
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
Ore-forming system
56 &?i = −0.5%₀
56 Fe (%₀ )
30.1%
Pyrite
1.00
0.00
−1.00
−2.00
−3.00
−4.00
−5.00
−6.00
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
62.5%
0.43‰
−0.33%₀
56 Fe (%₀ )
56 &?i = 0%₀
66 Zn (%₀ )
2.00
1.00
0.00
−1.00
−2.00
−3.00
−4.00
−5.00
−6.00
15
0.01
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
0.99
56 Fe (%₀ )
Geofluids
F
(l)
Figure 8: The Fe-Zn isotopic fractionation models for the ore-forming system and minerals (pyrite, sphalerite, and Mn-Fe carbonate) with
different 𝛿56 Fe𝑖 and 𝛿66 Zn𝑖 values.
[54], and 0.23‰ (the mean value of deep sea water) [63, 108,
109].
These fractionation models show that the 𝐹 ranges for
ore-forming system highly depend on the 𝛿𝑖 values (Figure 8).
The details are as follows: the pyrite covers the 𝐹 ranges
of 30.1%∼62.5% (𝛿56 Fe𝑖 = 0‰; Figure 8(a)), 48.8%∼99.0%
(𝛿56 Fe𝑖 = −0.5‰; Figure 8(b)), and more than 77.7% (𝛿56 Fe𝑖
= −1‰; Figure 8(c)). The 𝐹 ranges for sphalerite are
38.7%∼70.9% (𝛿66 Zn𝑖 = −0.28‰; Figure 8(d)), 26.7%∼48.7%
(𝛿66 Zn𝑖 = 0‰; Figure 8(e)), and 19.8%∼36.3% (𝛿66 Zn𝑖 =
0.23‰; Figure 8(f)), respectively. In comparison, the Mn-Fe
carbonates has the 𝐹 ranges of 7.1%∼24.9% (𝛿56 Fe𝑖 = 0‰;
Figure 8(g)), 10.9%∼39.7% (𝛿56 Fe𝑖 = −0.5‰; Figure 8(h)),
and 17.3%∼63.3% (𝛿56 Fe𝑖 = −1‰; Figure 8(i)) for Fe isotope,
as well as 55.6%∼68.1% (𝛿66 Zn𝑖 = −0.28‰; Figure 8(j)),
38.2%∼46.8% (𝛿66 Zn𝑖 = 0‰; Figure 8(k)), and 28.5%∼34.8%
(𝛿66 Zn𝑖 = 0.23‰; Figure 8(l)) for Zn isotope. All of these
results suggest that the Fe-Zn isotopic data of Zhaxikang
deposit fit these Rayleigh fractionation models well.
However, we need to consider the following facts in the
Zhaxikang deposit: (1) the second pulse of mineralization
has brought some Fe to form the stage 3 pyrite with heavier
𝛿56 Fe values (0.23‰∼0.43‰) [26]; thus the stage 3 pyrite
does not fit the fractionation models; (2) as most of the
sphalerite and pyrite are paragenetic during the first pulse of
mineralization (Figures 4(g), 4(i), 4(k), 5(a)–5(c) and 6(c)),
especially in the earliest lamina (Figures 4(a)–4(c), 5(c), 6(f)
and 6(m)), the sphalerite and pyrite should overlap on 𝐹
16
ranges; (3) in theory, the Mn-Fe carbonates would have the
same 𝐹 range for Fe-Zn isotopes. Nonetheless, in view of the
tight Zn isotopic variation range, the 𝐹 range for Fe isotope
should cover that for Zn isotope; (4) during the earlier period,
as the ore-forming system consisting of vapour and liquid
phases is dominant, 𝐹 values in fractionation models should
be large. Taking all these facts and 12 fractionation models
into consideration, the 𝛿56 Fe𝑖 value is supposed to be in
the range of −0.5‰∼−1‰, and the 𝛿66 Zn𝑖 value should be
between −0.28‰ and 0‰.
5.3. Implications for the Genesis of Zhaxikang Deposit
5.3.1. Excluding the Possibility of Hot Spring Genetic Model.
The hot spring model predicts that metals (e.g., Zn,
Pb, Sb, Ag, and Fe) are leached from sedimentary wall
rocks, which is supported by the following evidence: (1)
the 𝛿34 S values of the sulfides (4.5‰∼12‰) are similar to those of sedimentary wall rocks (4.93‰∼11.49‰);
(2) the 𝛿30 Si values of quartz (−0.90‰∼−0.40‰) are the
same as those of siliceous rocks with hot spring genesis;
(3) the 𝛿DV–SMOW (−135‰∼−127‰) and 𝛿18 OH2 O values
(−13.7‰∼12.4‰) of fluid inclusions trapped in quartz are
similar to those of the south Tibetan hot spring; (4) the
Pb isotopes (206 Pb/204 Pb: 18.474∼19.637; 207 Pb/204 Pb: 15.649∼
15.774; 208 Pb/204 Pb: 39.660∼40.010) show the characteristics
of radiogenic Pb; (5) the He-Ar isotopes demonstrate the
contribution of crustal fluid and meteoric water [3, 4].
However, this genetic model is inconsistent with textural
and Fe-Zn isotopic evidence presented here. Firstly, the
primary sedimentary wall rocks in the orefield are slate. As
suggested by the continuous batch experimental research of
Fernandez and Borrok [27], the ore-forming fluid would
preferentially leach out the heavy Zn isotopes. Likewise, Chen
et al. [110] have analyzed the Zn isotopic compositions of
samples from 8 hot springs, and the results show most of the
hot springs have relatively constant and heavier 𝛿66 Zn values
(approximately 0.42‰) than host rocks (−0.42‰ to 0.14‰).
Therefore, if the metallogenic elements are leached from the
slate by hot spring, there would be some sphalerite with
heavier 𝛿66 Zn values than these slates in Zhaxikang deposit.
Nevertheless, the 𝛿66 Zn values of the slate from Zhaxikang
orefield range from −0.23‰ to 0.10‰ that are similar to
sphalerite, especially the unmodified slate sample with the
𝛿66 Zn value of 0.10‰ that is even a little heavier than that of
sphalerite (−0.38‰ to 0.07‰; Figure 9(b)) [26]. As the slate
has heavier Zn isotopic compositions than those of host rocks
from Chen et al. [110], the 𝛿66 Zn values of the hot spring in
Zhaxikang orefield would even be heavier than 0.42‰, which
are much heavier than the 𝛿66 Zn𝑖 values (−0.28‰∼0‰) of
the fractionation models in Section 5.2. Secondly, Sharam
et al. [46] have measured the 𝛿56 Fe values (−0.59‰ to
−0.12‰) of hot springs in Juan de Fuca Ridge, which is much
heavier than the 𝛿56 Fe𝑖 value (−0.5‰∼−1‰) gained from
the fractionation models. Moreover, the marine fluids usually
have lighter Fe isotopic compositions (Figure 9(a)); thus the
hot springs in Tibet would have heavier 𝛿56 Fe values than
those of Sharam et al. [46].
Geofluids
The possibility of hot spring genesis for the first pulse
of mineralization event can be excluded by Fe-Zn isotopic
data. Evidence for the second pulse of mineralization is based
on the fact that the Fe-Zn isotope values do not follow
similar concentric patterns with the ore textures as seen
in sample ZXK-2. Meanwhile, the evidence from Si-H-O
isotopes demonstrates that the second pulse of mineralization
may be related to hot spring. Additionally, the Fe-Zn isotopic
data demonstrate that the sedimentary wall rocks have not
provided significant amounts of metals, although the S-Pb
isotopic data show that these wall rocks constitute some
contribution for S-Pb [4, 5], whereas the similar Zn isotopic
compositions of slate and sphalerite suggest that they share
the same Zn origin.
5.3.2. Inconsistency with the Magmatic-Hydrothermal Fluid
Genetic Models. There are two genetic models for the
magmatic-hydrothermal fluid genesis. In the first model,
Duan et al. [5] considered that the genesis of Zhaxikang
deposit relates to the mid-low temperature magma-related
hydrothermal activity and that the metallogenic elements
are mainly sourced from the mixing of basement and the
sedimentary wall rocks. The evidence is mainly from the ZnS-Pb isotopes that we mentioned in Section 1.
Duan et al. [5] have analyzed the 𝛿66 Zn values of the
sulfides (−0.25‰∼0.03‰) and basement rocks (0.05‰∼
0.21‰). The dominating sedimentary wall rocks in the
orefield are slate (𝛿66 Zn values: −0.23‰∼0.10‰) [26], which
is formed by the epimetamorphism of shale and sandstone.
Meanwhile, combining the data from Wang et al. [26] with
this study, the sphalerite should have a range from −0.38‰
to 0.07‰ in Zhaxikang deposit. Both the basement and
sedimentary wall rocks have heavier Zn isotopic compositions than sphalerite in Zhaxikang deposit. However, just
like we discussed in Section 5.3.1, if the Zn is sourced from
mixing of the basement and sedimentary wall rocks, there
should be some sphalerite with heavier 𝛿66 Zn values than
these rocks. This inference is also evidenced by the research
of Zhou et al. [64]: the Paleozoic carbonate host rocks and
Precambrian basements are considered to be the origin of
metals, and these rocks have lighter 𝛿66 Zn values (−0.52‰
to 0.16‰) than the sphalerite from the Tianqiao (−0.54‰ to
0.30‰) and Bangbangqiao (−0.21‰ to 0.43‰) deposits in
the Sichuan-Yunnan-Guizhou Pb-Zn metallogenic province
(Figure 9(b)). Additionally, in respect of the Fe isotope, the
Schwarzwald hydrothermal vein deposit in Germany can
be used as an analogy [52]. Iron in this deposit originates
from the basement consisting of granites and gneisses, as
well as sedimentary rocks including shale and sandstone.
The basement of Zhaxikang is composed of dolerite, quartz
diorite, rhyolite porphyry, pyroclastics, and porphyritic monzogranite [5], and these crust-derived igneous rocks have the
similar Fe isotopic composition with granites according to
the data from previous studies [21–23, 28, 36–45]. Likewise,
the 𝛿56 Fe values of shale and sandstone in the Schwarzwald
deposit are −0.21‰, 0.03‰, and 0.22‰, all of which fall into
the Fe isotopic variation range of shale (−0.39‰∼0.71‰)
from Beard et al. [34] and Rouxel et al. [35]. Therefore, we
Geofluids
17
56 F?i value
−1%₀∼−0.5%₀
66 ZHi value
−0.28%₀∼0%₀
Earth
0.03 ± 0.09%₀
Earth
0 ± 0.05%₀
Deep sea water: 0.23‰
Zhaxikang Sb-Pb-Zn-Ag deposit
Zhaxikang Sb-Pb-Zn-Ag deposit
Dongshengmiao SEDEX-type deposit, China
Dongshengmiao SEDEX-type deposit, China
Red Dog SEDEX-type ore district, America
Renison Sn-W deposit, Australia
Alexandrinka VHMS-type deposit, Russia
Schwarzwald hydrothermal
vein deposit, Germany
Cévennes MVT-type deposit, France
Irish-type
deposit, Ireland
Fenghuangshan
Skarn-type deposits in
Tongling ore district Dongguashan
China
Xinqiao
Basements
and ore-hosted
rocks
Bangbangqiao
Tianqiao
Iron ores in Bayan Obo Fe-REE
magmatic deposit,China
Carbonated-hosted Pb-Zn
Sulfide Deposit, Southwest China
Skarn-type deposits in Tongling ore district, China
Gorno and Raibl magmatic deposits, Italy
Midoceanic ridges pyrite
Seafloor hydrothermal
fluid system
Seafloor hydrothermal fluid system
Deep-sea carbonate
Igneous rock
Igneous rock
Shale
–2.5
–2.0
–1.5
–1.0
–0.5
0.0
0.5
1.0
Sedimentary rock
1.5
2.0
–0.8
–0.6
–0.4
–0.2
Pyrrhotite
Siderite
Chalcopyrite
Hematite
0.2
0.4
0.6
0.8
1.0
1.2
Zn (%₀ )
Fe (%₀ )
Pyrite
Mn-Fe carbonate
Arsenopyrite
Sphalerite
0.0
66
56
Magnetite
Bornite
(a)
Mn-Fe carbonate
Sphalerite
Slate
(b)
Figure 9: (a) Fe isotopic compositions of pyrite, Mn-Fe carbonate, and arsenopyrite in the Zhaxikang deposit (part of the data is quoted from
Wang et al. [26]). Other Fe isotopic data for the Bulk Silicate Earth [10], shale [34, 35], igneous rocks [21–23, 28, 36–45], seafloor hydrothermal
fluid system [34, 46–49], midoceanic ridges pyrite [48, 50], the Bayan Obo Fe-REE magmatic-type deposit in China [51], the skarn-type
deposits in Tongling ore district in China [21, 22], the Schwarzwald hydrothermal vein deposit in Germany [52], the Renison Sn-W deposit
in Australia [23], and the Dongshengmiao SEDEX-type deposit in China [53] are also plotted for comparison. (b) Zn isotopic compositions
of sphalerite, Mn-Fe carbonate, and slate in the Zhaxikang deposit (part of the data is quoted from Wang et al. [26]), compared with the Bulk
Silicate Earth [54], sedimentary rocks [55–57], igneous rocks [28, 54, 57–60], deep-sea carbonates [61], seafloor hydrothermal fluid system
[62], deep sea water [51, 54, 63], and other deposits with different geneses: the Gorno and Raibl magmatic-type deposit in Italy [14], the
skarn-type deposits in the Tongling ore district in China [12], the Tianqiao and Bangbangqiao carbonated-hosted Pb-Zn sulfide deposits in
China [64], the Irish-type deposit in Ireland [19], the Cévennes MVT deposit in France [65], the Alexandrinka VHMS-type deposit in Russia
[18], the Red Dog SEDEX-type ore district in Alaska [20], and the Dongshengmiao SEDEX-type deposit in China [53].
regard the fact that the metal-sourced rocks in these two
deposits have the similar Fe isotopic compositions. And, yet,
Markl et al. [52] suggested that fluid-rock interaction make
the ore-forming fluid have the 𝛿56 Fe value of −0.5‰∼0‰,
which is heavier than the 𝛿56 Fe𝑖 value (−0.5‰∼−1‰) of
the fractionation models in Section 5.2. Furthermore, it is
generally known that the equilibrium isotope fractionation is
a function of temperature, with larger fractionation generated
at lower temperatures [111]. Although the temperature of oreforming fluid (100∼200∘ C) [52] in the Schwarzwald deposit
is lower than that in Zhaxikang deposit (∼250∘ C) [5], the
lightest 𝛿56 Fe value (−1.36‰) of siderite from this deposit is
much heavier than that of Mn-Fe carbonate (−1.95‰) from
Zhaxikang deposit (Figure 9(a)). All of these inconsistencies
from Fe-Zn isotopic data indicate that this genetic model is
not appropriate for Zhaxikang deposit.
In the second genetic model, Xie et al. [6] suggested the
mineralization in Zhaxikang deposit was genetically related
to the Miocene dome-related magmatism. This magmatism generated the pegmatite and two-mica granite in the
core of the regional domes (Figure 1(b); e.g., Cuonadong,
Yalaxiangbo, Ranba, and Kangma), and the high-temperature
ore-forming fluids were derived from magmatic melts exsolution. This hypothesis is based on the research of field geology,
petrography, melt and fluid inclusions, and C−H−O isotopes: (1) the 𝛿13 CV–PDB (−6.1‰∼−6.9‰) and 𝛿18 OV–SMOW
(+9.9‰∼+11.8‰) values of rhodochrosite show the mantle origin; (2) the 𝛿13 DV–SMOW ( −144.8‰∼−110‰) and
𝛿18 OV–SMOW (−9.85‰∼−8.89‰) values, the low salinity
(0.2∼7.9 wt.% NaCleqv), high temperature (298∼457∘ C), and
rich CO2 content with minor CH4 , N2 , C2 H6 , C3 H8 , and
C6 H6 of the melt and fluid inclusions trapped in quartz and
beryl within pegmatite indicate the magmatic origin; (3) the
40
Ar-39 Ar plateau age of pegmatite (18.93 ± 0.27 Ma) is similar
to stage 5 quartz-pyrite-stibnite (17.9 ± 0.5 Ma).
The pegmatite and two-mica granite in the domes have
the typical characteristics of S-type granite: (1) containing
abundant Al-rich minerals, A/CNK: 1.07∼1.24; (2) w(SiO2 ):
73.26%∼74.33%, w(K2 O)/w(Na2 O): 1.1∼1.2, w(FeO)/w(Fe +
Mn): 0.64∼0.76; (3) the content of corundum molecules > 1%
in the CIPW standard minerals [112, 113]. Meanwhile, based
on the facts that all the Fe-bearing minerals in Zhaxikang
18
deposit are ferrous, and the magmatic-hydrothermal fluid
is CO2 -rich with minor CH4 , N2 , C2 H6 , C3 H8 , and C6 H6
[6], and the parent magma is most likely S-type reduced
magma. In consideration of the situation in the Renison SnW deposit that we mentioned in Section 1, the ore-forming
fluid is considered to exsolve from S-type reduced magma
and the sulfides have heavier Fe isotopic compositions
than ore-related igneous rocks [23]. The case in Zhaxikang
deposit is contrary yet the pyrite from the first pulse of
mineralization has the 𝛿56 Fe value of −0.33‰∼0.23‰ that
is much lighter than those of granitoids (−0.08% to 0.59%)
[22, 28]. Moreover, Heimann et al. [43] proposed that we
would not expect Fe isotopic compositions of high F/Cl magmatic/hydrothermal systems to significantly deviate from the
average of igneous rocks, hence as Xie et al. [6] considered
that the Zhaxikang parent magma has high F contents, the
magmatic-hydrothermal fluid in this genetic model would
have the similar Fe isotopic compositions with granitoids,
which is not in line with our fractionation models. Even
if the parent magma in Zhaxikang orefield is oxidized-type
and is similar to the magma related to I-type granitoids
with 𝛿56 Fe values of −0.04%∼0.59% in Tongling ore district
(Figure 9(a)) [21, 22], it is still hard to produce so light 𝛿56 Fe𝑖
value (−0.5‰∼−1‰). Similarity, Zn isotopic data do not
support this genetic model, either. Chen et al. [54] studied
the Zn isotopic fractionation during igneous process and
suggested that the maximum Zn isotopic variation induced
by high-temperature igneous processes is no larger than
𝛿66 Zn∼0.10%. Besides this, Telus et al. [28] measured the
𝛿66 Zn values of pegmatite (0.25‰∼0.59‰) and some other
granitoids (−0.16‰∼0.21‰) and then found there is even
no variation in 𝛿66 Zn values during fluid exsolution in some
cases. As the granitoids in the regional domes are principally
dominated by pegmatite and the magmatic fluids which have
high temperature of 298∼457∘ C [6], it is also hard to generate
𝛿66 Zn𝑖 value between −0.28‰ and 0‰ as yielded from the
Fe-Zn isotopic fractionation models. Consequently, the FeZn isotopic data are also not in favor of the second genetic
model.
On the other hand, neither of these two genetic models can explain the different Fe-Zn isotopic and elemental
variations in sample ZXK-2, which is considered to result
from the overprint by the second pulse of mineralization.
Overall, all of the Fe-Zn isotopic and elemental evidences
are inconsistent with both of these magmatic-hydrothermal
fluid genetic models. Nevertheless, the evidence for these
two genetic models may prove that the second pulse of
mineralization is related to magmatic-hydrothermal fluids.
5.3.3. Constraints on SEDEX Modified by Hydrothermal Fluid
Genetic Model. Zheng et al. [7, 25] considered that the first
pulse of mineralization (Pb-Zn) has the SEDEX genesis, and
the second pulse of mineralization (Sb-Ag) is related to hot
spring that overprints the earlier mineralization. The previous
evidences are mainly as follows: (1) the evidence from ore
textures and obviously late Sb mineralization compared to
Pb-Zn mineralization in Section 2.3; (2) the stage 1 and
2 ores have high Mn, Fe, Ba, and B contents, Ga ≫ In
Geofluids
(Ga/In: 1.49∼4.47), Pb + Zn ≫ Cu, and host rocks are
exhalative rocks and Mn-Fe carbonates; (3) the 𝛿34 S values
of sphalerite and galena (7.7‰∼12‰) are different from
those of stibnite (4.5‰200B∼7.1‰); (4) the data from valentinite (𝛿30 Si values: −0.9‰∼−0.4‰) and fluid inclusions
trapped in quartz (𝛿DV–SMOW : −162‰∼−142‰, 𝛿18 OH2 O :
−12.9‰∼−1.9‰, homogenization and freezing temperatures: 164∼313∘ C and −3.2∼−0.7∘ C, salinity: 0.7%∼5.3%, and
density: 0.74∼0.93 g/cm3 ) both demonstrate the Sb mineralization is related to the south Tibetan hot spring; (5) the
Rb-Sr isotopic isochrone age of stage 2 sphalerite is 147.2 ±
3.2 Ma that is similar to the Jurassic sedimentary wall rocks,
whereas a later stage quartz-pyrite-stibnite vein has the 40 Ar39
Ar plateau age of 17.9 ± 0.5 Ma.
Wen et al. [17] investigated several Pb-Zn deposits with
different geneses in China, and the results show that the
Zn/Cd ratios of sphalerite vary with different geneses: (1)
high-temperature systems including the porphyry, magmatic
hydrothermal, skarn, and volcanic hosted massive sulfide
(VMS)-type deposits: 155∼223; (2) low-temperature systems
that include the Mississippi Valley-type (MVT) deposits: 17∼
201; (3) SEDEX-type deposits of exhalative systems: 316∼
368; (4) seafloor hydrothermal sulfides of exhalative systems:
211∼510. According to the EPMA data, the Zn/Cd ratios of
sphalerite range from 296 to 399 in Zhaxikang deposit, which
overlap the range of exhalative systems and much higher than
those of high-temperature and low-temperature systems.
The Fe-Zn isotopic data also conform to the SEDEX
modified by hydrothermal fluid genetic model. Firstly, as
discussed in Sections 5.3.1 and 5.3.2, neither the hot spring
nor the magmatic-hydrothermal fluids can have the 𝛿56 Fe𝑖
(−0.5‰∼ − 1‰) and 𝛿66 Zn𝑖 (−0.28‰∼0‰) values to satisfy
the Fe-Zn isotopic fractionation models. However, the FeZn isotopic compositions of seafloor hydrothermal fluid
system covers the range of −1.79‰∼0.04‰ for Fe isotope
and −0.28‰∼0.96‰ for Zn isotope according to previous
studies (Figure 9) [38, 46–49, 62], which can generate the
ore-forming system with 𝛿56 Fe𝑖 and 𝛿66 Zn𝑖 values to meet
the fractionation models. Secondly, although there are large
overlaps in Zn isotopic compositions among deposits with
different geneses, the Zn isotopic compositions of Zhaxikang
deposit is most similar to the Alexandrinka VHMS-type
and Red Dog SEDEX-type deposits with marine origin,
as well as obviously distinguishing from the narrow range
of magmatic-hydrothermal deposits (Figure 9(b)). Thirdly,
the lightest 𝛿56 Fe value in Zhaxikang deposit is −1.95‰,
and only the minerals with marine origin (mid-oceanic
ridges pyrite) or SEDEX genesis (sphalerite and pyrrhotite
in Dongshengmiao SEDEX-type deposit) can have so light
𝛿56 Fe values (Figure 9(a)). Fourthly, under the conditions of
high 𝑃CO2 and low PH (<8), the Zn precipitated as sulfides is
isotopically nearly unfractionated with respect to the parent
hydrothermal fluid, whereas, under the conditions of high
𝑃CO2 and high-PH (>9), negative 𝛿66 Zn values down to 0.6‰
can be expected in sulfides precipitated from hydrothermal
fluid [107]. In Zhaxikang deposit, 𝑃CO2 would be high as
there are plenty of Mn-Fe carbonates; meanwhile, owing to
the facts that the modern seawater has the PH around 8
Geofluids
and there is more CO2 in Jurassic atmosphere than present,
the Jurassic seawater would have a lower PH than modern
seawater. This can well explain that the 𝛿66 Zn values of
sphalerite (−0.38‰∼0.07‰) slightly fractionate with the
ore-forming system (𝛿66 Zn𝑖 value: −0.28‰∼0‰). Besides
these, the overprinting of earlier ores by second pulse of
mineralization is not only proved by the different Fe-Zn
isotopic and elemental variations in sample ZXK-2 from
the other 3 samples in this research (Figure 5) but also
evidenced by the temporally increasing 𝛿56 Fe and decreasing
𝛿66 Zn values recorded in this deposit that coincided with an
increase in alteration [26]. Nevertheless, further research is
required to confirm whether this hydrothermal fluid is hot
spring or magmatic-hydrothermal fluid.
To sum up, among the various genetic models, the Fe-Zn
isotopic and EPMA evidence indicate the SEDEX modified by
hydrothermal fluid genetic model is the most plausible. Our
research also demonstrates that the Fe-Zn isotopes have the
potential to trace the metal source and provide insights into
ore-forming processes.
6. Conclusions
The EPMA and Fe-Zn isotopic data allow us to make the
following conclusions:
(1) The ore-forming elements are transported by the
ore-forming system that is the mixture of vapour
and liquid phases; the vapour-liquid partitioning and
mineral precipitation are the main cause of Fe-Zn
isotopic and elemental variations.
(2) The Fe-Zn isotopic fractionation models demonstrate that the 𝛿56 Fe𝑖 and 𝛿66 Zn𝑖 values of the oreforming system are in the range of −0.5‰∼−1‰ and
−0.28‰∼0‰, respectively.
(3) Based on the evidence from the EPMA data, FeZn isotopic characteristics, and fractionation models,
the SEDEX modified by hydrothermal fluid genetic
model is most plausible for the Zhaxikang deposit.
(4) There are two pulses of mineralization in the Zhaxikang deposit; the overprint by the second pulse of
mineralization has also partly modified the Fe-Zn
isotopic and elemental compositions of some earlier
samples.
Conflicts of Interest
The authors declare that they do not have any commercial
or associative interest that represents conflicts of interest in
connection with the submitted work.
Acknowledgments
This work was carried out while Da Wang was a visiting
student at the Juniata College. The authors would like to
thank Matthew Gonzalez (Pennsylvania State University) and
Linda Godfrey (Rutgers University) for aid in measuring
and access to the Neptune instruments. They acknowledge
19
support from the Program for Changjiang Scholars and
Innovative University Research Teams (IRT14R54, IRT1083),
the Commonwealth Project from the Ministry of Land and
Resources (201511015), and the Fundamental Research Funds
for the Central Universities (2652015044 and 2652015354).
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