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Consistency between apatite and zircon
petrochronology supports robustness of apatite in
ngerprinting igneous processes in porphyry
systems
Hongying Qu
Institute of Mineral Resources, CAGS
Julie Rowland
University of Auckland
Jingwen Mao
Institute of Mineral Resources, CAGS
Evan Orovan
British Columbia Geological Survey
Michael Rowe
University of Auckland
Shihua Zhong
Ocean University of China
Research Article
Keywords: Robustness of apatite, zircon, petrochronology, in-situ multi-element analysis, porphyry–skarn
system
Posted Date: June 24th, 2024
DOI: https://doi.org/10.21203/rs.3.rs-4524703/v1
License: This work is licensed under a Creative Commons Attribution 4.0 International License.
Read Full License
Additional Declarations: No competing interests reported.
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Abstract
Apatite low-temperature thermochronology can be double or even triple dated allowing for a
reconstruction of the thermal history of rock from ~ 550 oC to near-surface temperatures. Even though it
has disadvantageous U–Th–Pb contents (high Pb contents and low U and Th contents) and an unstable
nature, apatite is still regarded to have the same robustness in ngerprinting igneous processes in
porphyry systems as zircon, so far as to be replace zircon. Hence, we systematically studied
characteristics of morphology, geochronology and geochemistry of apatite hosted in syenogranite and
monzogranite intrusive rocks in the large Hutouya skarn deposit, in order to corroborate its potential
thermochronological monitoring capabilities like zircon in ngerprinting igneous processes in porphyry
systems. In this study, apatite grains can be subdivided into two types, FI-free Apatite I formed in the
early less fractionated magma and FI-rich Apatite II crystallized in the late highly fractionated magma
stage. We obtained ages of 229.0 ± 6.6 Ma in syenogranite and 224.3 ± 4.5 Ma / 223.7 ± 3.9 Ma in
monzogranite from Apatite I of magmatic origins. Zircon grains in the two granites can be classied into
three types. Zircon I is characterized by transparent and bright zones, Zircon II by dark and metamict
features, and Zircon III by mineral inclusions. Zircon I grains with a magmatic texture of well-developed
bright oscillatory zones, are most likely primary magmatic zircon that crystallized early in the evolution of
granitic magma, dating results of which are 224.70 ± 0.61 Ma in syenogranite intrusions and 225.75 ±
0.66 Ma / 226.31 ± 0.78 Ma in monzogranite, respectively. The apatite–zircon timing is coincident.
Furthermore, apatite trace rare earth element contents in the syenogranite and monzogranite intrusions
display a negative-slope chondrite-normalized distribution from La to Lu with strong negative Eu
anomalies and weak positive Ce anomalies, with major element contents that are statistically identical
with enriched F but poor Cl. Zircon trace element compositions in the two intrusions show consistent
and steeply increasing chondrite-normalized REE diagrams from La to Lu with negative Eu anomalies
and strong positive Ce anomalies. Accordingly, apatite U–Pb dates and the corresponding in-situ trace
element compositions and isotopes can test precise constraints on rock formation ages, temperature,
oxygen fugacity, material source, and tectonic background, which can be relatively more robust when
used as proxies for magma oxidation state.
1. Introduction
The determination of rock-forming and mineralizing ages is critical to understanding mineral deposit
genesis. Zircon, as is stable, has a high sealing temperature, and high U and low common Pb
concentrations, making it the most commonly used imeral for U–Pb dating; however, zircon may be
absence, or may not represent the timing of mineralization in some mineral deposits. Recently,
developed highly sensitive analytical techniques, most notably, laser ablation inductively coupled plasma
mass spectrometry (LA–ICP–MS) age dating of U-rich minerals (e.g., apatite, epidote, cassiterite,
magnetite, garnet, titanite, rutile, calcite, wolframite, scheelite, etc.), has allowed researchers to not solely
rely on zircon for dating hydrothermal events in mineral deposits (Andersson et al., 2019; Glorie et al.,
2020; Mao et al., 2016; Pan et al., 2016; Qu et al., 2019b). It is worth mentioning that apatite, being a
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common accessory mineral in magmatic rocks and hydrothermal deposits, is generally employed in low-
temperature thermochronology research (Chew and Spikings, 2015; Prowatke and Klemme, 2006;
Webster et al., 2009). Due to its relatively low U–Th–Pb isotope sealing temperature (~ 550–350 oC),
apatite dating can provide age information for P–T–t trajectory research for metamorphic rocks with
complex evolutionary and thermal histories. Apatite can be double or even triple dated (U–Pb, ssion
track, and U–Th–Sm/He), allowing a reconstruction of a thermal history of a rock from ~ 550 oC to near-
surface temperatures (Carrapa et al., 2009; Glorie et al., 2019; Jepson et al., 2018). Although apatite LA–
ICP–MS dating has advantages, such as in-situ rapid age determinations, there are signicant problems
for samples with high Pb and low U and Th contents. Also, apatite is unstable in acidic groundwaters and
weathering proles and has only limited mechanical stability in sedimentary transport systems (Morton
and Hallsworth, 1999). In contrast, zircon have a much more stable nature than apatite. Like zircon,
apatite trace element compositions can be used to interpret characteristics of the melt from which it
was derived, e.g., its compositional evolution, degree of assimilation and fractionation, oxidation state,
and even determination of paragenetic separation from the parental magma (Ballard et al., 2002; Liang et
al., 2006; Odlum et al., 2022; Zhang et al., 2017). This is possible because the apatite mineral structure,
Ca5(PO4)3(F,Cl,OH), can incorporate a variety of transition metals, rare earth elements (REE), and other
cations. For example, Sr2+, Mn2+, Fe2+, REE, and Na+ can be substituted in the Ca2+ and Si4+ sites, and
As5+ and S6+ can be substituted in the P5+ site (Hughes and Rakovan, 2002).
In this study, we evaluate the consistency between apatite and zircon petrochronology from rocks that
are proximal to a skarn deposit to test the robustness of apatite in ngerprinting igneous processes in
porphyry systems. We used LA–ICP–MS to obtain U–Pb ages of apatite and zircon for Middle–Late
Triassic intrusions of the Hutouya skarn deposit, NW China, and determined their in-situ trace element
compositions. Through comparisons with global trace element compositions of apatite from fertile
intrusions in the porphyry systems, we concluded that a similar robustness in ngerprinting igneous
processes in porphyry systems as zircon, and that apatite thermochronology can test complementary
information on magma evolutions and ore-forming uids that are not available from zircon alone.
2. Regional geological setting
China has one of the widest distribution of skarn deposits in the world, providing for the domestic
industrial demand of 70% tin, 60% tungsten, 30% copper, 23% molybdenum, 20% gold, and 11% iron
(Zhao et al., 2013). However, distributions of skarn deposits in China are irregular, with the vast majority
(over 95%) of large and medium-sized deposits occurring in East China (i.e., Pacic Rim metallogenic
domain), especially Fe–Cu skarns, which occur in the middle and lower reaches of the Yangtze River and
the Yanliao Cu–Fe–Mo polymetallic mineralization belt. In recent years, there has been an increased
discovery rate of large- and medium-sized polymetallic skarn deposits in the Qiman Tagh Metallogenic
Belt (QMB) of the East Kunlun Mountains, forming the most prole polymetallic skarn belt in the
northwest region (Feng et al., 2011, 2010).
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The QMB is in the western portion of the East Kunlun Orogenic Belt (EKOB) along the northern part of the
Qinghai–Tibet Plateau (QTP), between the Qaidam Basin and Kumukuri Basin in NW China (Feng et al.,
2010). It is an important exploration target area for porphyry- and skarn-related Fe–Cu–Pb–Zn deposits
(Feng et al., 2012; Zhao et al., 2013; Zhong et al., 2021). At present, skarn-related Fe–Cu–Pb–Zn
deposits are the main prospecting targets in the belt, and many economically-mineable skarn deposits
have been discovered, including the Kendekeke Fe (Xiao et al., 2013), Hutouya Cu–Pb–Zn–Fe (Feng et
al., 2011; Qu et al., 2019a), Kaerqueka Cu (Wang et al., 2009), Galinge Fe (Yu et al., 2015), and Yemaquan
Fe deposits (Gao et al., 2014).
The QMB is associated with the Qiman Tagh Orogen, which was constructed through protracted
accretion and collision of a collage of terranes during the subduction and closure of the Qiman Tagh
Ocean, a branch of the Paleo-Tethys Ocean from the Neoproterozoic to Early Mesozoic (Wang et al.,
2009; Yu et al., 2017). The early Neoproterozoic (ca. 1000–820 Ma) ages for this orogen suggests a link
with the formation of the Rodinia supercontinent (He et al., 2016; Meng et al., 2015). The Qiman Tagh
Terrane was tectonically and chronologically separated into the North Qiman Tagh Terrane (NQT) and
South Qiman Tagh Terrane (SQT), which was tectonically clipped by the Adatan fault in the east and
Baiganhu fault in the west (Wang et al., 2009; Yu et al., 2017) (Fig.1A). The NQT was an active
continental margin containing abundant Paleozoic granitoids, which possibly formed through melting of
old basement (Li et al., 2013; Wang et al., 2014). In contrast, the SQT was an exotic terrane that had intra-
oceanic subduction, where supra-subduction zone (SSZ) type ophiolites were documented together with
island arc tholeiite and calc–alkaline lavas, in a primary oceanic island arc environment during the Early
Paleozoic (Meng et al., 2015). In addition, the SQT developed abundant Late Paleozoic and Early
Mesozoic granitoids (Chen et al., 2006; Yao et al., 2016) (Fig.1B, C). The collision between the SQT and
NQT occurred probably in the Late Silurian (ca. 422 Ma) and continued to ca. 398 Ma (Chen et al., 2006;
Yao et al., 2016), as evidenced by ages of abundant within-plate granitic magmatism in the NQT that
formed after 398 Ma (Yao et al., 2016). The nal closure of the Paleo Tethyan Qiman Tagh Ocean might
have occurred in the Late Permian, and resulted in the accretion of the Kumukuri microcontinent
followed by signicant Triassic magmatism (Chen et al., 2005; Feng et al., 2012). A series of calc–
alkaline and alkaline granitoids generated through mantle–crustal mixing were linked with transitions
from post-collision to within-plate settings (Yu et al., 2015).
3. Geology of the Hutouya skarn
The Hutouya skarn, located at the center of the QMB (Fig.1D), hosts a Cu–Pb–Zn resource of
0.85million tones (Mt) at an average grade of 2.05% Cu, 5.79% Pb, and 4.46% Zn and an Fe resource of
200 Mt at an average grade of 28.82% Fe (Zhao et al., 2013). It comprises both magnesian and calcic
skarns carrying Fe–Sn–Cu–Co mineralization in the inner zone and Pb–Zn mineralization in the outer
zone, and locally contains W–Mo–Ag–Bi–Sn mineralization, and substantial pyrrhotite-bearing iron ores.
Skarn alteration and Fe–Cu–Pb–Zn–W–Mo–Ag–Co–Bi–Sn mineralization is developed at the contacts
between carbonate rocks and granitic intrusions. The ore is typically hosted by E–W trending Lower
Carboniferous Dagangou marble and limestone and Upper Carboniferous Di’aosu Formation marble,
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calcsilicate hornfels, and dolomitic limestone with thinly-bedded limestone (Fig.2). Mesoproterozoic
Langyashan Formation of the Jixian Group, Ordovician–Silurian Qiman Tagh Group, and Upper Triassic
Elashan Formation are also exposed in the district. All Paleozoic sedimentary rocks in the Hutouya
deposit are extensively faulted and folded along E–W trends related to compressive fracture zones.
These structures appear to be important in ore localization. More detailed descriptions can be found in
Qu et al. (2019a).
Indosinian Permo–Triassic intermediate to felsic intrusions in the Hutouya deposit area include red
syenogranite, light buff-colored monzogranite, and gray granodiorite, and diorite (Fig.3). The
syenogranite and monzogranite intrusions are spatially associated with mineralization in Ore Belt III
(Fig.4A) and show an intrusive contact relationship (Fig.4B). The syenogranite is widespread in the
center of the ore district and occurs as a 1.4 km2 stock intruding the Di'aosu Formation, the Qiman Tagh
Group, and Dagangou in the northern, western, and southern part of the district, respectively (Fig.2). The
monzogranite only crops out sporadically in the south of the Ore Belt II (Fig.2). The irregularly-shaped
monzogranite stock has an exposed area of 9 km2, and contains large amounts of mac microgranular
enclaves (MMEs) (Fig.4C, D). The intrusion has a contact with the underlying Di'aosu Formation where
thick magnetite-rich skarn formed.
The Hutouya deposit includes 51 skarn ore bodies in seven ore belts. These include three Cu–Pb–Zn ore
bodies, and several medium-sized Fe ore bodies locally associated with W–Mo–Ag–Bi–Sn
mineralization (Fig.2).
Ore belts I, II, and III contain Cu and Mo as an “inner contact” skarn with minor Fe and Sn. These skarns
developed on and near contacts of syenogranite and monzogranite with carbonate-rich strata of the
Lower Carboniferous Dagangou and Upper Carboniferous Di'aosu Formations. At the inner contact zone
with the intrusions, the syenogranite displays strong endoskarn, alteration, and radial and meshed skarn
veins (Fig.4E). However, the monzogranite shows a sharp contact relationship with the skarn and
syenogranite intrusions in Ore belt III, which indicates that the monzogranite was probably emplaced
after the syenogranite and skarn alteration (Fig.4F). Some of the meshed skarn veins occurr in the
surrounding marble and are interpreted as the metasomatized front. The geometry and extent of these
ore bodies are controlled by structures along the intrusive contact zone. Chalcopyrite, pyrite, and
magnetite hosted in banded and massive skarns are the dominant ore minerals in these ore belts (Liu et
al., 2013), with lesser pyrrhotite, arsenopyrite, and minor stannite (Zou et al., 2017).
Ore belts IV–VII host Pb–Zn “outer contact” skarns with subordinate copper mineralization. The Pb–Zn
ore bodies occor mainly in fracture zones and are largely strata bound sphalerite- and galena-bearing
skarn bodies selectively replacing carbonate layers of the Di'aosu Formation, Qiman Tagh Group, and
Langyashan Formation. Chalcopyrite and magnetite also occur in those ores. The ore belts and ore
bodies appear to be both structurally- and stratigraphically-controlled.
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Skarn bodies range a few meters to tens of meters in width, and discontinuously extend to more than
two km in length. Metal contents of skarns in the deposit are zoned from innermost Mo-bearing ores
nearest intrusive contacts progressively outward to Fe–Sn–Cu–Co, Cu–Mo–(Pb–Zn), and outermost
Pb–Zn zones. In a broad sense, silicate mineralogy in the skarn zones from innermost garnet-rich skarn
near potassic-altered igneous rock grades progressively outward to diopside-, epidote-, tremolite- /
actinolite-, and chlorite-rich minerals, surrounded by a peripheral zone of recrystallized marble with local
massive suldes. This zoning is complicated by late-stage quartz–sulde and phlogopite-rich retrograde
alteration that crosscuts early prograde skarn. Liu et al. (2013) suggested that uids responsible for
retrograde alteration played an important role in concentrating sphalerite in the Pb–Zn ores.
4. Sampling and analytical methods
4.1. Sampling
There three representative rock samples from the Hutouya deposit were selected for petrographic
examination, as well as U–Pb in-situ trace element analysis. Sample HT1901 was taken from the quartz-
rich syenogranite and Sample HT1903 (which was divided into two subsamples) was selected from the
porphyritic monzogranite. The samples are all from the Cu–Mo “inner contact” zone of Ore belt III.
Zircon grains were separated using magnetic and heavy liquid separation. Approximately, 1000 zircon
grains from each sample were mounted and polished in 25-mm epoxy discs. The 400 apatite grains used
in this study were not separated by standard mineral separation techniques; rather, they were selected by
employing optical and back-scattered electron (BSE) microscopy from polished thin sections. This
approach combines detailed textural relationships, allowing for a more precise interpretation of apatite
compositions.
Individual apatite and zircon grains show conspicuous euhedral to subhedral columnar shapes in BSE
(Fig.5) and cathodoluminescence (CL) (Fig.6), and there are no obvious cracks on the surface.
4.2. Analytical methods
4.2.1. BSE and CL imaging
BSE imaging of apatite was performed by electron-probe microanalysis (EPMA) at the MNR Key
Laboratory of Metallogeny and Mineral Assessment, using a JEOL JXA-8800 instrument with a 2 to 5 µm
beam.
Internal zonation patterns of zircon crystals were observed in CL images at the MNR Key Laboratory of
Metallogeny and Mineral Assessment, Institute of Mineral Resources, Chinese Academy of Geological
Sciences (CAGS), Beijing, China. Using a combination of CL imaging and optical microscopy, the clearest
and least fractured zircon crystals were selected as suitable targets for laser ablation.
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4.2.2. U–Pb dating and in-situ multi-element analysis of
apatite and zircon
U–Pb dating and in-situ multi-element composition analysis of apatite and zircon was performed using
an ASI RESOLution S-155 ablation system with Coherent Compex Pro 110 Ar-F excimer laser operating at
a 193 nm wavelength and pulse width of 20ns coupled to an Agilent 7900 quadrupole ICP–MS. Detailed
analytical conditions are described in Thompson et al. (2016). All instrumentation is housed at the
Centre for Ore Deposit and Earth Sciences (CODES) Analytical Laboratory at the University of Tasmania
(UTAS), Hobart, Australia.
Analyses comprise a 30s blank gas measurement followed by a further 30s of ablation when the laser is
switched on using a 29 µm spot size, ring at a frequency of 5 Hz and beam energy density of 2.0 J/cm2
for zircon and 3.5 J/cm2 for apatite. All analyses have a pre-ablation of 5 laser shots to remove any
surface contamination. Ablation was performed in a pure He atmosphere owing at 0.35 L/min and
immediately mixed with Ar, owing at a rate of 1.05 L/min after ablation. The mass of each isotope (e.g.,
23Na, 31P, 43Ca, 51V, 56Fe, 88Sr, etc.) was measured every ~ 2 ms with longer counting times on the Pb and
U isotopes.
Data reduction in apatite was done using the method outlined in Thompson et al (2016) and references
therein, where a common Pb correction was performed on the calibration standard. The downhole
fractionation, instrument drift and mass bias correction factors for Pb/U ratios were calculated using
values of the OD306 apatite from Thompson et al (2016). The calibration of the U–Pb ages was
monitored using several apatite reference materials: 401 apatite (Thompson et al., 2016), Durango
apatite (McDowell et al. 2005) and the McClure Mountain apatite (Schoene and Bowring 2006). Ages are
calculated using Concordia intercept with Stacey and Kramer’s (Stacey and Kramer, 1975) model Pb
composition at the age of the apatite, unless there was enough spread on the isochron to negate the
need for the assumption of common Pb composition.
In zircon, U–Pb dating was based on the method outlined in Thompson et al (2018) and Halpin et al.
(2014). For each analysis, a subset of data that closely matches a concordant composition was selected
for quantication. The downhole fractionation, instrument drift, and mass bias correction factors for
Pb/U ratios were calculated from analyses of the 91500-zircon using the values of Wiendenbeck et al.
(1995). A calibration of U–Pb ages was performed by comparing measured analyses of the Temora
zircon (Black et al. 2004) and the Plesovice zircon (Slama et al. 2008) with published values. Trace
element abundances measured in the 91500 zircon were within the range of reported values from the
GeoReM website (http://georem.mpch-mainz.gwdg.de/). All common Pb corrections were done using
Stacey and Kramer’s (Stacey and Kramer, 1975) model Pb composition at the age of the zircon, unless
independent common Pb compositions existed for a sample.
In both routines, instrument drift and mass bias correction factors of the 207Pb/206Pb ratio (ages) were
determined using the Pb isotopic values of the NIST610 glass determined by Baker et al. (2004). Trace
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element abundances were calibrated on the NIST610 glass from values of Jochum et al. (2011) and
using secondary standard corrections based on the composition of the glasses BCR-2G and GSD-1G
(GeoReM preferred values). Quantication was performed using 43Ca as an internal standard element in
apatite and 91Zr in zircon and normalizing all measured cations to stoichiometric concentrations of
these elements in each respective mineral. The calibration standards and the NIST610, BCR-2G and GSD-
1G glasses were analyzed in duplicate at the beginning, end and every 60 minutes throughout the
analytical session. All data was processed using the program LADR (v1.1.01; Norris and Danyushevsky,
2018).
4.2.3. Apatite major element analysis
Major element analysis of apatite was performed using EPMA at the UTAS using a JEOL JXA-8800
instrument with a 2 to 5 µm beam. F, Na, and Cl were analyzed with a 4 nA beam current and 10 kV
accelerating voltage in the rst instrumental pass; the remaining elements were measured utilizing a 20
nA beam current and 20 kV accelerating potential in the second instrumental pass. Natural minerals and
synthetic oxides were used as standards, and the ZAF software provided by JEOL was used to correct
matrix effects. The accuracy of the analytic results is 1–5% depending on the abundance of the element.
5. Results
5.1. U–Pb ages of apatite and zircon
Typical BSE images of apatite from both intrusives are shown in Fig.5. Age data of apatite are
summarized in Table S1 and Fig.7. Apatite from both intrusives can be categorized into two types: FI-
free Apatite I and FI-rich Apatite II (Fig.8). In this study, we present dating results of Apatite I. Most
apatite from the two intrusives are euhedral, elongate, and tabular crystals about 100 × 40 µm to 200 ×
60 µm in size without prominent zonation patterns (Fig.5). The apatite age of sample HT1901
(syenogranite) yielded a lower intercept age of 229.0 ± 6.6 Ma, (MSWD = 1.15, n = 51), and samples
HT1903-1 and HT1903-2 (monzogranite) yielded lower intercept ages of 224.3 ± 4.5 Ma (MSWD = 3.5, n
= 27) and 223.7 ± 3.9 Ma (MSWD = 1.4, n = 40) (Fig.7).
Typical CL images of zircon are shown in Fig.6, and U–Pb data are summarized in Table S2 and Fig.9.
According to size, color, texture, and morphology, zircon crystals from the two intrusives can be
classied into Zircon I with transparent and bright zones, Zircon II with dark and metamict features, and
Zircon III with mineral inclusions (Fig.8). In this study, we present dating results of Zircon I. The zircon
grains are clear to pale, euhedral to subhedral, nearly granular (normally 100 to 200 µm in size) and
display concentric zonation patterns in CL. Zircon U–Pb ages of the two intrusives show a grouping with
lower intercept ages in Tera–Wasserburg Concordia diagrams. The zircon age from sample HT1901
(syenogranite) yielded a lower intercept age of 224.70 ± 0.61 Ma (MSWD = 1.3, n = 24) (Fig.9), and
samples HT1903-1 and HT1903-2 (monzogranite) yielded lower intercept ages of 225.75 ± 0.66 Ma
(MSWD = 1.15, n = 23) and 226.31 ± 0.78 Ma (MSWD = 1.4, n = 16) (Fig.9), respectively.
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U–Pb ages of apatite–zircon from the two intrusives (monzogranite and syenogranite) in the Hutouya
deposit are consistent within error. These ages likely represent the time of emplacement of the
intrusives.
5.2. In-situ trace element compositions of zircon and
apatite
Trace element contents of apatite (Apatite I) are summarized in Table S3 and Fig.10. V, Pb, Th, and U
concentrations of the apatite range from − 1 ppm to 10s ppm and do not show systematic variations
from between the two intrusive. In contrast, Sr contents are much higher, ranging from 10s ppm to 100s
ppm. REE concentrations range from 1,000s ppm to over 1 wt% in both intrusives. In the chondrite-
normalized REE diagrams (Fig.10), results have a negative slope from La to Lu, strong negative Eu
anomalies with similar (Eu/Eu*)N, and weak positive Ce anomalies with similar (Ce/Ce*)N values
(Fig.11).
Trace element compositions of zircon are listed in Table S4. Syenogranite (HT1901) and monzogranite
(HT1903-1 and HT1903-2) have average REE contents of 732 ppm, 756 ppm, and 992 ppm, respectively.
Th and U contents range from 10s ppm to 1,000s ppm. Hf concentrations are 1,000s ppm, with average
values of 10,195 ppm for syenogranite, and 9,771 ppm and 10,069 ppm for the monzogranite samples.
Chondrite-normalized REE diagrams for zircon show a consistent and steeply increasing trend diagrams
from La to Lu with strongly positive Ce anomalies and negative Eu anomalies (Fig.10). According to the
(Ce/Ce*)N calculation method proposed by CODES on the basis of a lattice-strain model for mineral-melt
partitioning of Ce4+ and Ce3+ cations, the relationships of Ce and Eu anomalies among different
intrusions at Hutouya are examined. Calculated zircon (Ce/Ce*)N and (Eu/Eu*)N values are listed in Table
S4, where the subscript indicates chondrite normalization. Zircon all display higher (Ce/Ce*)N values,
ranging from 6 to 421 (average: 119) in the syenogranite and 15 to 309 (average: 84) and 1 to 245
(average: 62) in the monzogranite samples. The (Eu/Eu*)N values do not show noticeable differences
between the intrusives, with most values ranging from 0.05 to 0.52 (Fig.12).
5.3. Apatite major element compositions
Analytical data of major elements for apatite from syenogranite (HT1901) and monzogranite (HT1903-1
and HT1903-2) are summarized in Table S5. The consistency of the results and previous published
studies for the secondary standards suggests that the major element compositions for apatite are
robust and reliable. The low analytical totals in some of the analyses are likely related to: 1) OH not being
analysed, 2) elements present in the mineral but not captured in the routine, 3) the laser crater is
adjacent to the EPMA spot in a suciently small grain, such that the interaction volume either overlaps
with the crater or with epoxy, 4) charging effects due to insucient coating of the crater next to the
EPMA spot, and 5) there could be beam damage,since apatite are very susceptive to it.
Major element contents of apatite from syenogranite and monzogranite intrusives are statistically
identical. All apatite samples are uorapatite, enriched in F but have low Cl contents.
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6. Discussion
6.1. Robustness of apatite petrochronology in ngerprinting
porphyry systems
Accurate chronological constraints are critical for establishing the timing of mineralization, deciphering
mineralization processes, and developing mineral exploration models. Geochronological and
geochemical ngerprints of mineralization processes can be preserved by apatite and zircon. Zircon has
characteristics of high stability and sealing temperatures, and high U and low common Pb contents, and
therefore is one of the most suitable minerals for U–Pb dating. However, zircon may be absent in some
ore systems or may not directly represent mineralization. Therefore, in recent years, the development of
LA–ICP–MS dating of other U-rich minerals that form in hydrothermal uids or related intrusives has
become a powerful method to determine the age of mineral deposits. Apatite, a common accessory
mineral in magmatic rocks and hydrothermal deposits, stands out due to it already being widely
employed in low temperature thermochronology research (Chew and Spikings, 2015; Prowatke and
Klemme, 2006; Webster et al., 2009). Apatite can be U–Pb dated (Tc = 550–350 oC), ssion track dated
(Tc = 110–60 oC), and U–Th–Sm/He dated (Tc = 80–40 oC), forming a medium- to low-temperature
continuous thermochronology that can comprehensively and continuously analyze tectono–
thermochronological history (Carrapa et al., 2009; Glorie et al., 2019; Jepson et al., 2018). Furthermore,
apatite can accommodate a variety of elements (e.g., S, Sr, U, Th, REE, etc.) and has high volatile
contents, such as F, OH and Cl, making it an ideal mineral for both geological dating and tracing (Chu et
al., 2009; Piccoli and Candela, 1994).
In this study, we selected apatite and zircon grains from the syenogranite and monzogranite in the
Hutouya deposit to study thermochronology and in-situ trace element compositions, with the aim to test
the consistency between apatite and zircon for petrochronology and ngerprinting of igneous processes
in a porphyry–skarn system. Apatite grains in the two intrusives can be categorized into two types (FI-
free Apatite I and FI-rich Apatite II; Fig.8). The variations of textures and geochemical compositions in
the two types of apatite are indicative of changing crystallization environments. The euhedral grains of
Apatite I might be crystallized from a volatile-undersaturated magma. Conversely, Apatite II with lower Cl
and higher F contents are only distributed in the more highly fractionated granite (Table S5, Fig.8) and
formed under a volatile-oversaturated stage demonstrated by the rich uid inclusion contents
(Andersson et al., 2019; Glorie et al., 2020; Mao et al., 2016; Pan et al., 2016; Qu et al., 2019b). Lower Cl
and higher F contents of Apatite II could be attributed to the segregation of isolated uid phase in the
late aqueous magma (Chu et al., 2009; Doherty et al., 2014; Mathez and Webster, 2005; Sha and
Chappell, 1999; Webster et al., 2009). Zircon grains in the two granites can be classied into Zircon I with
transparent and bright zones, Zircon II with dark and metamict features, and Zircon III with mineral
inclusions (Fig.8) in the CL images, indicating that they were formed under different physicochemical
conditions during the magmatic-hydrothermal evolution. Zircon I grains have a magmatic texture of well-
developed bright oscillatory zones, and are most likely primary magmatic zircon that crystallized early in
Page 11/32
the evolution of granitic magma. The low Th and U contents and higher Zr/Hf ratios of Zircon I (Table S4)
indicate they crystallized from volatileundersaturated anhydrous magma (Erdmann et al., 2013; Qu et al.,
2019b; Zeng et al., 2016). Zircon II occurring as individual grains or overgrowth with the Zircon I might be
of a successive later origin than the Zircon I (Fig.8). Notably, high Th and U contents of Zircon II may be
its crystallization from a volatile-enriched aqueous magma (Nasdala et al., 2001; Claiborne et al., 2006;
Bacon et al., 2007; Geisler et al., 2007; Erdmann et al., 2013). Zircon III grains full of numerous
hydrothermal mineral inclusions might be of the product of uid interaction with previous Zircon II in a
volatile-oversaturated environment, indicative of hydrothermal crystallization or hydrothermal alteration
(Breiter and Skoda, 2012; Erdmann et al., 2013; Hoskin, 2005; Hoskin and Schaltegger, 2003). Collectively,
apatite and zircon from the syenogranite and monzogranite in the Hutouya deposit experienced a
prolonged crystallization process and were altered by late-stage exsolved uids. The well-developed
euhedral apatite and oscillatory primary magmatic zircon represent an early-crystallized phase from a
least fractionated granite. Metamict zircon occurs as individual grains or overgrowth with the magmatic
zircon formed under volatile–saturated aqueous magma during the magmatic-hydrothermal transition
stage. Some formed zircon was altered by exsolved magmatic uids in the most fractionated granite,
indicating a volatile oversaturated environment. Meanwhile, apatite with abundant uid inclusions and
high F/Cl ratios from the most fractionated granite crystallized in this subsolidus stage.
Apatite I and Zircon I are interpreted to be of magmatic origins, and their ages therefore represent the
time of magma emplacement. The apatite LA–ICP–MS U–Pb ages range from 235–220 Ma, which is
consistent with the zircon U–Pb ages that range from 227–224 Ma. These ages are indistinguishable
within error and indicate that the magmas were emplaced over a short time span, and cooled rapidly,
given the closure temperature of apatite (~ 620 oC) and zircon (900 oC).
As hydrothermal activities and mineralization in porphyry–skarn systems are intimately tied to the
emplacement of ore-forming intrusions (Razique et al. 2014), the associated hydrothermal and
mineralization events at Hutouya probably have similar short durations of just a few million years or less
(Zhong et al., 2018). At the regional scale, magmatism at Hutouya coincides with that in the other
porphyry and skarn deposits in the QMB (Fig.1C). For instance, Kaerqueka porphyry–skarn in the QMB
formed circa 227 Ma (Feng et al., 2012), 224 Ma at Yazigou Cu–Mo deposits (Li et al., 2008), and 229.4
Ma at Kendekeke Fe deposits (Xiao et al., 2013). The granitoids associated with skarns in the area
formed during the same period, including the ages at Hutouya (this study), 225 Ma at Galinge Fe skarn
deposits (Zhao et al., 2013), and 227 Ma at Tawenchahan Fe skarn deposits (Feng et al., 2012).
It should be noted that geochemical compositions of apatite can be regarded as a tool to identify
magmatic mineralization potentials. The magmatic oxygen fugacity is a key factor to fertile magmas of
porphyry deposits (Lehmann, 1990; Sun et al., 2015, 2013; Wittenbrink et al., 2009; Zhang et al., 2017;
Zhong et al., 2017). Oxidized magmas are more likely to form Cu porphyry deposits than reduced
magmas (Imai, 2002; Li et al., 2017a; Liang et al., 2006; Lu et al., 2016), considering that under the high
oxygen fugacity, sulfur in magmas mainly exists in the form of sulfate (SO42−) which has a much higher
solubility in silicate melts than sulde, that is, sulde is dicult to reach saturation then precipitate
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during the magmatic stage thus facilitating metal accumulations in the late stage of magmatic
evolutions (Ballard et al., 2002; Richards, 2003). Skarn and porphyry deposits have similar magmatic
origins and evolution processes (Li et al., 2017b), so it is also applicable to Cu (–Pb–Zn) skarn deposits
in controlling of the oxygen fugacity to fertil magmas of Cu porphyry deposits.
Zircon (Eu/Eu*)N and (Ce/Ce*)N values are effective indicators for evaluating the magmatic oxygen
fugacity (Ballard et al., 2002; Gardiner et al., 2017; Trail et al., 2012). In the Hutouya Fe–Cu–Pb–Zn skarn
deposit, the syenogranite and monzogranite have similar (Eu/Eu*)N and (Ce/Ce*)N values in zircon, (Table
S4; Fig.12). According to the Weibao Cu–Pb–Zn skarn deposit in the QMB, fertile intrusions have higher
Ce4+/Ce3+ values than those of non-fertile intrusions (Zhong et al., 2018). We can draw a conclusion that
parental magmas with the higher oxygen fugacity from fertile intrusions in Cu skarn deposits tend to
form Cu mineralization. However, Pb and Zn are not easily controlled by oxygen fugacity and behave as
incompatible elements. This means that magmas related to the Pb–Zn mineralization can be either high
oxygen fugacity or low oxygen fugacity. Many studies have shown that both S-type granite (low oxygen
fugacity magmas) and I-type granite (high oxygen fugacity magmas) can form Pb–Zn skarn deposits (Fu
et al., 2017; Niu et al., 2017), which also supports that Pb-Zn mineralization is independent of magmatic
oxygen fugacity conditions. In other words, the oxidation state is not a controlling factor for Pb–Zn
mineralization within the Hutouya deposit.
Apatite Eu and Ce anomalies may be more easily affected by other factors unlike zircon Eu and Ce
anomalies which are mainly controlled by oxygen fugacity conditions, therefore the relationship between
apatite Eu and Ce anomalies and magmatic oxygen fugacity is not so similar (Piccoli and Candela,
1994). Nevertheless, if the physical conditions (specically temperature and pressure) and
concentrations of these elements in magma are relatively stable, apatites crystallizing from more
oxidized magma will have higher Eu3+/Eu2+ but lower Ce4+/Ce3+ than reduced magma owing to ion
substitution in the apatite structure, which results in apatites having strong negative Eu and positive Ce
anomalies (Cao et al., 2012; Sha and Chappell, 1999). In this study, we further conrm that magmatism
in the Hutouya deposit is similar to other skarn deposits in the QMB. Notwithstanding, this work shows
that the fertile intrusions at Hutouya can be well dened by (Eu/Eu*)N and (Ce/Ce*)N parameters, that is
strong negative Eu anomalies and weak positive Ce in apatite. Hence, apatite Ce anomalies including
(Ce/Ce*)N, Ce4+/Ce3+, and Ce/Nd values are relatively more robust as proxies for magma oxidation state
(Loader et al., 2017). The parameter (Eu/Eu*)N, although affected by many magmatic processes, can still
reect the magma redox state to some degree (Dilles et al., 2015). Moreover, previous studies found that
Mn contents are signicantly controlled by the oxygen fugacity with high apatite Mn contents in reduced
magmas while low apatite Mn contents in oxidized magmas, which can be explained by the substitution
of Ca2+ by Mn2+. Compared with Mn3+ and Mn4+, Mn2+ is more easily enriched in apatite, because the
ionic radius of Mn2+is close to that of Ca2+ (Belousova et al., 2002). Mn mainly exists as Mn2+ at low
oxygen fugacity with high Mn contents in apatite, while Mn mainly exists as Mn3+ and Mn4+ at the high
oxygen fugacity with low Mn contents. In the Weibao deposit, apatite within two ore-forming intrusions
Page 13/32
exhibits much lower Mn concentrations than within the barren diorite porphyry. Apatite Mn contents of
the two fertile intrusions in the Hutouya deposit are similar to the fertile intrusions in the Webao deposit
(Zhong et al., 2018).
Notwithstanding, halogen contents of magmas (especially F and Cl contents) in apatite can also be
regarded as an important indicator to evaluate productive magmas since halogens can effectively
complex and transport metal elements (Coulson et al., 2001; Pan et al. 2016; Piccoli and Candela, 1994;
Webster et al., 2004). A previous study showed that apatite, occurring as inclusions within biotite and
hornblende, was one of the early crystal phases that appeared during crystallization (Tang et al., 2021),
and thus the F and Cl partitioning between the apatite and the melt seems unlikely to be inuenced by
the crystallization of biotite and hornblende. Therefore, the contents of chlorine and uorine in the melt
predominantly affected by their magmatic sources can be evaluated from the concentrations of chlorine
and uorine in apatite. In this study, the uorine contents of apatite is much higher than chlorine
contents, because the partition coecient of uorine between apatite and melt is much higher than
those of chlorine (Mathez and Webster, 2005). In addition, the chlorine contents of the two fertile
intrusions in this study remain almost invariable, because the apatite/melt ratio is approximately
constant at low Cl contents, but when the melt becomes saturated in Cl both partition coecients
increase rapidly as Cl content of the bulk system increases (Doherty et al., 2014). Furthermore, magmas
formed by partial melting of lower crust materials usually show relatively stronger enrichment of F and
depletion of Cl than those formed by dehydration melting in slab subduction environments (Ding et al.,
2015; Jiang et al., 2018; Kendrick et al., 2011; Xu et al., 2022). Therefore, the volatile components of the
parent magmas of syenogranite and monzogranite are mainly related to lower crustal melting.
6.2. Implication for regional exploration
The QMB region experienced two important tectonic evolutionary processes of the Proto-Tethys Ocean
and the Paleo-Tethys Ocean, corresponding to two magmatic cycles of the Early Paleozoic and Late
Paleozoic to Early Mesozoic (Feng et al., 2012; Mo et al., 2007; Yang et al., 2003; Yu et al., 2017). The
Proto-Tethys Ocean began to form and expand in the Early Cambrian (Feng et al., 2010; Yang et al.,
1996), the subduction gradually weakened in the Silurian and began to transition to a collisional orogeny
stage (Liu et al., 2013; Ren et al., 2009), and changed from a syn-collision compressional environment to
a post-collision extensional environment in the Devonian (Meng et al., 2015; Qi et al., 2016; Zhao et al.,
2008). The Paleo-Tethys Ocean was in a subduction stage during the Late Permian-Early Triassic (Yao et
al., 2020), and entered collision and post-collision stages during the Middle–Late Triassic (Feng et al.,
2012). The Middle–Late Triassic is a very important metallogenic period in the QMB (Gao et al., 2014;
Wang et al., 2018; Zhang et al., 2010). At this stage, the QMB evolution changed from a compressional
and transpressional to an extensional environment, which resulted in asthenosphere upwelling and
strong crust–mantle interaction (Yao et al., 2017). As consequence, extensive partial melting of lower
crust caused widespread development of magmatic intrusions in the upper crust. Thus, it provided
favorable conditions for polymetallic mineralization in this area, and ore-forming ages of skarn
Page 14/32
polymetallic deposits are concentrated in a range of 230–224 Ma (Qu et al., 2023), consistent with the
apatite and zircon aged determinations in this study.
7. Conclusion
Apatite from the two fertile intrusives (syenogranite and monzogranite) in the Hutouya skarn deposit can
be divided into FI-free Apatite I and FI-rich Apatite II, meanwhile, zircon can be classied into three sub-
types: Zircon I with transparent and bright zones, Zircon II with dark and metamict features, and Zircon III
with mineral inclusions.
They syenogranite and monzogranite were emplaced between 235 to 220 Ma (apatite ages), which is
similar to their zircon ages (227 to 224 Ma). These ages are coincident with other fertile intrusives in
QMB. These ages are coincident with other fertile intrusives in the QMB. They also have similar
magmatic oxygen fugacity coecients and apatite halogen contents as other fertile intrusives in the
QMB, indicating that apatite trace element compositions can be used as robust proxies for magma
oxidation state in porphyry-skarn systems.
Declarations
CRediT authorship contribution statement
Hongying Qu: data curation, funding acquisition, investigation, writing original draft, and writing review
and editing. Julie Rowland: supervision and writing review and editing. Jingwen Mao: conceptualization,
supervision, and writing review and editing. Evan Orovan: data curation, methodology, and writing review
and editing. Michael Rowe: formal analysis and methodology. Shihua Zhong: conceptualization and
formal analysis
Declaration of competing interest
The authors declare that they have no known competing nancial interests or personal relationships that
could have appeared to inuence the work reported in this paper.
Acknowledgments
This study was nancially supported by the China Geological Survey Program (Grant number:
DD20240117). Prof. Chengyou Feng and Dr. Miaoyu, Hui Wang, and Jiannan Liu are acknowledged for
their assistance during the eldwork. We are thankful for assistance from Prof. David Cooke and Leonid
Danyushevsky and Dr. Lejun Zhang for LA–ICP–MS analyses.
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Figures
Figure 1
(A) Tectonic outline of the Tibet Plateau showing the locations of the main suture zones. The satellite
image of the East Kunlun showing major tectonic units. (B) Tectonic–magmatic sketch map of the
Qiman Tagh showing tectonic subdivisions, South Qiman Tagh magma belt and North Qiman Tagh
magma belt. They are bordered by the Adatan thrust fault in the east and Baiganhu thrust fault in the
west. (C) Frequency plot of the U–Pb ages for the intrusions from the North Qiman Tagh Belt and South
Qiman Tagh Belt. (D) The age data sources are referred in the text. Ages that are in ~200 Ma are colored
blue, ~300 Ma are black, and ~400 Ma are red. Italic characters: SHRIMP; standard character: LA–ICP–
MS.
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Figure 2
Geological sketch map of the Hutouya deposit in Qiman Tagh, Qinghai Province (modied after Feng et
al., 2011), showing sample locations in this study (triangles).
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Figure 3
Micrograph of syenogranite and monzogranite (cross-polarized light).
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Figure 4
Field photographs of Ore Belt III in the Hutouya skarn deposit. (A) Contact relationship between Ore Belt
III and intrusions; (B) intrusive relationship between monzogranite and syenogranite; (C, D) MMEs hosted
by monzogranite; (E) endoskarn showing strong alteration of syenogranite; (F) monzogranite intruded
into altered syenogranite and skarn.
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Figure 5
Representative BSE images of apatite for syenogranite and monzogranite in the Hutouya deposit. The
analytical spots and dating results are shown.
Figure 6
Representative CL images of zircon for syenogranite and monzogranite in the Hutouya deposit. The
analytical spots and dating results are shown.
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Figure 7
Diagrams of apatite U–Pb dating for syenogranite and monzogranite in the Hutouya deposit.
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Figure 8
A schematic illustration of the zircon and apatite recording the magmatic–hydrothermal evolution
process (modied after Qu et al., 2019b). The variations of the zircon types (I, II, and III) and apatite types
(I and II) on morphology, textures and geochemical compositions reecting the progressive evolution of
the granitic melt from volatile-undersaturated to oversaturated conditions.
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Figure 9
1) Diagrams of zircon U–Pb dating for syenogranite and monzogranite in the Hutouya deposit, and 2)
histograms on ages and numbers.
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Figure 10
Chondrite-normalized REE diagrams of zircon and apatite for syenogranite and monzogranite in the
Hutouya deposit.
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Figure 11
Apatite (Eu/Eu*)N versus (Ce/Ce*)N diagram for syenogranite and monzogranite in the Hutouya deposit.
Figure 12