The transition to new and green technologies is increasing the need for metals1,2 such as copper for which demand is forecast to increase by 140–350% from 2010 to 20503,4. Porphyry-type deposits provide more than 70% of global copper, around 95% of molybdenum and important amounts of gold (20%) and other metals5. Most form from hydrothermal fluids associated with large and long-lived calc-alkaline to slightly alkaline, water-rich and relatively oxidising trans-crustal magmatic systems, mainly in subduction-related settings e.g.5,6,7,8,9. Whilst such systems are arguably rather common, porphyry-, and particularly large porphyry-type deposits are extremely rare and increasingly difficult to find7. Their formation may require a series of specific conditions and events during the evolution of magmatic-hydrothermal systems.
In the drive to discover new ore deposits, there have been many recent attempts to develop whole-rock and mineral geochemical indicators to assess whether certain magmatic systems may be significantly mineralised, or ‘fertile’10. Their main advantage compared with conventional exploration techniques is that they are relatively cheap and are of low environmental impact. Most indicators reflect the hydrous nature of the magmas from which porphyry-type deposits form e.g.7,11,12,13,14,15,16,17,18,19,20.
The current paradigm is that the hydrous magmas that form porphyry-deposits result from a long (multi-million year), arc-scale, subduction-driven ‘ramp-up’ in volatiles and ore-forming constituents in mid- to lower crustal magmatic reservoirs9,11,16,21,22,23,24. Before emplacement into the upper crust, the magmatic system develops its ore-forming geochemical signatures over protracted time scales, in excess of ~ 5 Myrs, due to cyclical fractionation and re-charge of deep reservoirs by mafic magmas. The subduction-related tectonic regime has been suggested to progressively deepen the melt evolution zone and/or slow the upwards migration of magmas through the crust9,22,25,26. Alternatively, the ore-forming potential of magmas, and associated geochemical signatures, may increase during evolution within an upper crustal staging ground e.g.5,27,28,29. Distinguishing the nature of magmatic evolution in the lead up to porphyry copper ore formation is problematic because of the paucity of vertically extensive exposure over the crustal windows of porphyry ore-forming systems30; this has resulted in a fragmented understanding of the magmatic timescales associated with porphyry-deposit formation.
To address this, the Yerington magmatic system, western Nevada, was studied as it has provided constraints for many of the most commonly used porphyry system models, mostly due to its unique ~ 8 km deep profile, from volcanic to plutonic environments, through at least four porphyry copper deposits (Figs. 1, 2) e.g.5,6,27,28,30,31,32,33. Here we reconstruct the Yerington magmatic system across the deep plutonic to volcanic environment, encompassing deep-seated melt evolution zones, through to the development and focusing of magmatic-hydrothermal fluids to form porphyry-type deposits. We present a new 4-D model based on the timescales and drivers for the evolution of the magmatic system’s ore-forming potential and associated geochemical signatures.
The construction of a porphyry-forming magmatic system
The middle Jurassic composite Yerington batholith27,34,35,36 lies within a volcanic-arc terrane, in the early Mesozoic marine province37, which formed due to subduction tectonics to the west35. The batholith was emplaced into Triassic to Jurassic intermediate composition volcanics, volcaniclastic and argillaceous sedimentary rocks, and basal exposures of the likely semi co-eval Jurassic Artesia Lake Volcanics, which are unconformably overlain by the Jurassic Fulstone Spring Volcanics (subaerial quartz-latitic to dioritic lavas, domes, ignimbrites and volcaniclastics)27,35,38 (Figs. 1, 2). Late Cenozoic extensional faulting and associated fault block rotation in the Basin and Range has exposed a < 1 to ~ 8 km palaeodepth cross-section through the Yerington batholith27,34,39 (Figs. 1, 2).
There are three main plutonic phases, which, listed in order of increasing emplacement depth, are: (1) the McLeod Hill quartz monzodiorite (McLeod QMD); (2) Bear quartz monzonite (Bear QM); and (3) Luhr Hill granite (LHG)27. These are cross-cut by swarms of granite-composition porphyry and aplite dykes27,33. Units of the Fulstone Volcanics are thought to have been cogenetic with granite porphyry dykes rooted in the LHG38,40, or, alternatively, may have been cogenetic with the nearby younger Shamrock batholith and post-date porphyry mineralisation35. The dyke swarms are spatially and temporally associated with the batholith’s four known porphyry copper deposits: Ann Mason; Yerington; MacArthur and Bear (Figs. 1, 2) e.g.27,41,42. Combined, these host a resource in excess of 9 Mt of contained Cu42,43,44,45.
Temporal constraints from field relationships
Field-based observations place constraints on the relative timing of magmatism, alteration and mineralisation. The LHG is the youngest of the three main plutons having been emplaced into the McLeod QMD and Bear QM27. Contacts between the LHG and previously emplaced plutons are sharp (Fig. S1), with no chilled margins or evidence of interaction with precursor granitoids. No metasomatic effects are present at the contacts beyond the later, pervasive, mostly sodic-calcic and propylitic porphyry-related alteration46,47. In deeper portions of the LHG (~ 7.5 km palaeo-depth, based on structural reconstructions27), banding is observed locally, defined by grain size variations (Fig. S2).
The onset of porphyry mineralisation is constrained by cross-cutting relationships; it is spatially and temporally associated with multiple generations of variably mineralised granite-composition porphyry and aplite dykes that clearly cross-cut the upper (Fig. 3; S3) as well as lower parts of the LHG, and appear to have been focused through apophyses of the LHG33,38,40,41,42,47,48. The dykes generally have sharp contacts with the LHG, with some showing chilled margins and others lobate contacts (Fig. 3a). It was previously suggested that both the porphyry and aplite dykes emanated from cupolas and upper zones (~ 3 to 6 km depth) of the LHG27, however we could not trace either to their source and therefore suggest that they were likely to have been intruded from below the deepest levels exposed in the LHG (> ~ 7 km). Different generations of aplite dykes either cross-cut and/or mingle with the porphyry dykes (Fig. 3b–d), which indicates multiple intrusion events, with some generations emplaced penecontemporaneously with porphyry dykes and others later. In the palaeo-vertically deepest (> 6 km) exposures of the LHG (or the ‘root zone’ for the porphyry deposits30), the majority of aplite dykes pre-date the spatially associated late-stage coarse muscovite veins and alteration48,49, as well as Na-Ca alteration. However, in deep exposures of the LHG, certain aplite dykes, which appear to post-date the muscovite and Na-Ca alteration, are thought to have been emplaced from a larger, longer-lived, deeper source48; as these post-date the hydrothermal alteration they are not considered further in this paper.
Despite the close temporal relationship between the porphyry and aplite dykes, they have very different textures. The porphyry dykes show no direct textural evidence for fluid exsolution (e.g. miarolitic cavities50), and are only seen to be cross-cut by mineralised veins. In contrast, certain generations of aplite dykes contain miarolitic cavities, pegmatitic segregations, early ‘A-type’ quartz ± chalcopyrite ± bornite ± molybdenite veins (nomenclature after51), quartz unidirectional solidification textures (USTs) which grow inwards from their margins, and are cross-cut by mineralised veins (Fig. 3c–f & S3-S6)52. These observations are comparable with previous descriptions of aplitic ‘vein dykes’ in other porphyry systems e.g.53,54,55. The presence of quartz USTs within the aplites is likely to indicate undercooling56 and rapid pressure fluctuations due to repeated carapace fracturing55,57 (which may induce fluid exsolution via first-type boiling58), suggesting that these mineralising aplite dykes were emplaced rapidly to shallow depths. Given that the aplite dykes host mineralised miarolitic cavities that are closely associated with early A-type mineralised veins (Fig. 3d–f & S4), they capture the nature and timing of magmatic-hydrothermal fluid exsolution and mineralisation. Here, we only focus on the generations of aplite dykes which cross-cut the LHG cupola and are directly associated with mineralisation.
Field relations indicate that some parts of the Fulstone volcanics were cogenetic with the emplacement of porphyry dykes associated with the LHG38,40. Propylitic alteration (e.g. epidote replacing primary plagioclase, and chlorite replacing mafic minerals; Fig. S7) is ubiquitous across the Fulstone Spring Volcanics, indicating that the hydrothermal system could have been active for some time after volcanism. The lack of more acid alteration (e.g. advanced argillic) may indicate that these volcanics, if related, were deposited away from the central axis of the porphyry system.
Absolute age constraints on magmatic system evolution
The determination of crystallisation ages for igneous samples using U–Pb CA–ID–TIMS on zircons (See Methodology and Supplementary Data 1) provides a temporal framework for the construction of the Yerington batholith and eruption of overlying volcanics, over an indicated period of ~ 2.8 Myrs (~ 169.3 Ma to ~ 166.1 Ma; Zircon ages are reported in Fig. 4 and Supplementary Data 2. A sensitivity analysis of the ages is presented in Supplementary Data 2). This supersedes the previous U–Pb geochronological framework which was based on multi-grain TIMS35, and ion probe analyses36 on limited sample sets and which had relatively large uncertainties.
According to the new schema, the McLeod QMD pluton (AM72 and AM4QMD) was emplaced over a period of > ~ 0.9 Myrs (~ 168.6 Ma to ~ 167.4 Ma) and the Bear QM pluton (BH10 and MA9) > ~ 1.5 Myrs (~ 169.6 Ma to ~ 167.1 Ma). The youngest zircons for these two units are indistinguishable in crystallisation age, indicating a period of contemporaneous emplacement with crystallisation of their latest phases within ~ 100 kyrs of each other. These mineralogically and texturally distinct plutons were likely emplaced episodically to form their internal contacts27,34 and both appear to young downwards over a palaeo-vertical distance of ~ 3 km, supporting under-accretion e.g.59 as the mode of emplacement32.
Within the LHG, its upper region (AM18LHG; ~ 5 km palaeo-depth27) and a deeper portion (LH24S; ~ 7.5 km palaeo-depth27) show closely comparable zircon 206Pb/238U dates and weighted means of 167.365 ± 0.041 Ma and 167.275 ± 0.027 Ma, respectively (Fig. 4). These ages define the maximum emplacement time-gap between the McLeod QMD-Bear QM and LHG of 215 ± 59 kyrs. However, when the 167.440 ± 0.039 Ma age of a mineralised porphyry dyke from within the Ann Mason porphyry deposit (AC25) is considered, which cross-cuts the LHG cupola ~ 1 km higher in the system than AM18LHG, then the maximum emplacement time-gap must be shorter (140 ± 57 kyrs, or ~ 100–200 kyrs). These cross-cutting relationships imply an episodic emplacement of the exposed LHG over > ~ 150 kyrs and the new timescale that construction of the Yerington batholith was at least two times longer than the ~ 1 Myrs previously estimated on the basis of previous geochronology32,35.
A stratigraphically lower unit of the Fulstone volcanics (BS1) yielded an age (168.318 ± 0.054 Ma; Fig. 4) within the emplacement duration defined by the Bear QM and McLeod QMD, whereas the stratigraphically higher unit (BS16) gave a much younger age (166.285 ± 0.059 Ma; Fig. 4), ~ 1.1 Myrs younger than the formation of the LHG cupola. This indicates the volcanic record spans over ~ 2 Myrs, rather than there having been a single post-ore volcanic event, as proposed by35.
From cross-cutting relations, the onset of porphyry-style Cu-Mo mineralisation in Yerington is temporally constrained by the emplacement of dyke swarms through the cupolas of the LHG e.g.27,41,42. In turn, our absolute U–Pb ages for the mineralised porphyry dykes that cross-cut the cupolas of the LHG, and the youngest ages for the McLeod QMD and Bear QM, both constrain the onset of ore formation to ~ 167.4 Ma (Fig. 4). The multiple generations of aplite dykes that host mineralised miarolitic cavities and early A-type veins (AM63A and AM13BAP; Fig. 3d–f & S3-S6), and have been proposed to act as conduits for the transport of mineralising fluids into the ore-forming environment33, capture the timing, albeit a partial record, of magmatic-hydrothermal fluid exsolution and mineralisation. As the youngest zircon growth within these aplite dykes likely crystallised as part of the magmatic assemblage at the magmatic-hydrothermal transition (See QEMSCAN, Fig. S4–S6, Supplementary Data 3), the U–Pb ages of 167.282 ± 0.040 Ma and 167.045 ± 0.057 Ma (Fig. 4) constrain the timing of mineralisation to a period of at least ~ 400 kyrs.
From Re-Os molybdenite ages for chalcopyrite-bornite-molybdenite-quartz veins (A- and B-type) (samples AC11, AC12 & AC21), a chalcopyrite-molybdenite-bearing quartz UST within an aplite dyke (or vein dyke texture e.g.53,54,55) (AC3) and a fine grained molybdenite vein (AC41MP) (Fig. 4 & S8; Supplementary Data 2) from the Ann Mason deposit, mineralisation occurred during multiple hydrothermal events over a period in excess of 1.5 Myrs, from 166.90 ± 0.1 to 165.29 ± 0.1 Ma. Comparison between the hydrothermal Re-Os molybdenite ages and magmatic zircon U–Pb ages requires that systematic uncertainties, relating to the tracer calibrations and decay constant intercalibration, must be considered, which typically equate to ± 0.8 Myrs on Re-Os dates and ~ ± 0.08 Myrs for U–Pb (Refer to Supplementary Data 2). Results therefore indicate some component of porphyry-style mineralisation within the Ann Mason deposit could potentially have occurred coincident with the eruption of the younger propylitically altered components of the Fulstone volcanics (BS16), at 166.285 ± 0.059 Ma (Fig. S7). In general, the results indicate that hydrothermal mineralisation was not a single, short-lived event.
Geochemical change within the magmatic system
In terms of their whole-rock geochemical compositions, the McLeod QMD and Bear QM (pre-mineralisation) are similar and notably different to the LHG, porphyry and aplite dykes (syn-mineralisation) (Figs. 5, 6, S9 & S10; Supplementary Data 4). The McLeod QMD and Bear QM have similar ranges in SiO2 (~ 60–68 wt%) whilst the LHG samples either overlap with these or are marginally more evolved (~ 67–69 wt.% SiO2). The porphyry dykes show a range in SiO2 (60–71 wt%), whilst aplite dykes are the most evolved, generally having > 73 wt.% SiO2. Compared to the McLeod QMD and Bear QM, the LHG and porphyry dykes have higher Sr/Y ratios (Sr/Y > 130), steeper LREE/HREE and MREE/HREE patterns (e.g. La/Yb > 30; Gd/Yb > 3.7), lower ƩREEs (< 100 ppm), and positive Eu anomalies (Eu/Eu* > 1.05). Whilst the Dy/Yb values (~ 2) do not significantly change between the McLeod QMD, Bear QM and LHG, they follow a slightly negative trend with increasing SiO2 (Fig. S11).
From whole-rock geochemistry, the pre-mineralisation McLeod QMD and Bear QM are likely to be genetically related, despite their mineralogical and textural differences27. The more evolved composition of the syn-mineralisation LHG was probably due to a change in the bulk fractionating assemblage of the magmas. As previously shown27, this is likely to have been from clinopyroxene- (in which Y, MREEs and HREEs are compatible, although more weakly when compared to amphibole60,61) and plagioclase-dominated fractionation (in which Sr and Eu are compatible62) in the pre-mineralisation units, to deeper and wetter e.g.12,13 amphibole-dominated fractionation, with plagioclase crystallisation delayed until after emplacement into the upper crust, in the syn-mineralised units. The elevated melt-water contents led to higher Sr/Y and Eu/Eu* and depletion in HREEs60,61,63 (Figs. 5, 6). From the work of64, the slightly negative trend of Dy/Yb with increasing SiO2 in the plutonic units (Fig. S11) is likely to indicate that garnet did not play a role in the geochemical evolution of the system.
Zircon geochemistry is a function of pressure, temperature and melt composition65,66 and therefore records changes in the geochemical and physical nature of the melt from its source to level of emplacement, although only during the period of zircon saturation. Zircon from across the Yerington magmatic system (Fig. 7 & S12–S15; Supplementary Data 5) can be separated into two distinct geological groups: pre-mineralisation (McLeod QMD, Bear QM and older volcanic units) and syn-mineralisation (LHG, aplite dykes and younger volcanic units). Zircon Hf concentrations, typically thought to reflect melt evolution65, are comparable between the pre- and syn-mineralisation units. Zircon from the pre-mineralisation McLeod QMD and Bear QM have relatively higher Ti (5–20 ppm) and lower Eu/Eu* (0.2–0.5) and Gd/Yb (MREE/HREE, 8–21) compared with the syn-mineralisation intrusives, overlapping with the pre-mineralisation Artesia volcanics and older units of the overlying Fulstone volcanics. From outer to inner portions of the McLeod QMD and Bear QM, there is an increase in zircon Gd/Yb (rising from ~ 10 to ~ 16) with decreasing Ti concentration. In contrast, zircon from the syn-mineralised LHG, aplite dykes and younger units of the Fulstone volcanics have lower Ti (2–5 ppm), higher Eu/Eu* (~ 0.4 to 0.9) and Gd/Yb (~ 10–35).
There is no major difference in zircon composition between the LHG and aplite dykes. Further, the ‘early’ and ‘late’ mineralised porphyry dykes are almost identical in their zircon compositions, from both the Ann Mason and Yerington porphyry copper deposits, in agreement with previous zircon data for dykes from the Yerington porphyry deposit36. Regardless of age, the porphyry dykes show no clear division between the geochemical groups of pre- and syn-mineralisation plutonic rocks, which we attribute to recycling of pre-mineralisation zircon grains from a magmatic reservoir at depth. Similarities in zircon geochemical signatures between the intrusive units and the volcanics likely indicates that are genetically linked.
As melt chemistry is largely linked to the composition and processes in the evolution zone e.g.67, major differences in zircon chemistry, and by extrapolation melt chemistry, between the samples is likely to reflect differences prior to magma emplacement. The comparable zircon chemistry of the McLeod QMD and Bear QM indicate that these plutons had a shared source and evolution prior to emplacement, controlled by clinopyroxene- and plagioclase-dominated fractionation, despite their mineralogical and textural differences27. In contrast, the syn-mineralised units (LHG and aplite dykes) underwent amphibole-dominated fractionation, evidenced by increasing MREE/HREE, and suppressed plagioclase crystallisation in the source, caused by relatively high melt-water contents, prior to significant plagioclase crystallisation post emplacement into the upper crust, the latter indicated by relatively high Eu/Eu* values63. Elevated Eu/Eu* in zircon may also relate to increased melt fO2 e.g.8,29,68, but ∆FMQ values (calculated by the method of69) overlap between the pre- and syn-mineralising intrusives (Fig. S16), and zircon Eu/Eu* is not a robust proxy for melt redox conditions as it is also strongly controlled by the crystallisation of other phases within the melt70. The low Ti concentration seen in the inter-mineralisation LHG and aplite dyke zircons could reflect lower temperatures defined by the Ti-in-zircon geothermometer71, induced by increased melt-water contents and reduced temperatures at which zircon dominantly crystallises within the melt72. However, given the paucity of good constraints on the titania activity through the evolution of the magmatic system, calculation of absolute temperatures has been avoided. Low zircon Ti concentration within later magmatic phases could also be due to decreased titania activity in the magma, due to greater incorporation of Ti into amphibole, titanite, or other Ti-bearing phases crystallising at depth. Importantly, the changes indicative of a shift from a clinopyroxene-plagioclase-dominated system to an increasingly hydrous, amphibole-dominated system at the transition from a non-mineralising to mineralising magmatic system, both suggested in previous work27, and whole-rock data here (Figs. 5, 6), are more pronounced within the zircon geochemistry than whole-rock geochemistry, both in the plutonic and volcanic record.
Isotopic constraints on magma pathways
From whole-rock 87Sr/86Srt data27, there appears to have been a subtle transition from more radiogenic, crustal-like values73, within the pre-mineralisation intrusions, to less radiogenic ratios within the syn-mineralisation intrusions, which suggests a decreasing amount of crustal assimilation over time. Similarly, zircon O-isotopes also show a transition from δ18O ~ 6 ‰36, above values for zircon equilibrated with mantle-derived melts and indicating contamination with other crustal components, to ~ 4.5 ‰ and is within uncertainty of the expected range for the mantle74,75 (Fig. 8).
The ɛHft compositions of the zircon crystals that yield the youngest dates by CA–ID–TIMS U–Pb are a good approximation for the late-stage melt at the emplacement level and provide further insights into the evolution of the Yerington magmatic system (Fig. 8 & S17; Supplementary Data 6). Over the period of construction of the Yerington batholith, there were changes in the nature of zircon Lu–Hf isotopes. This is best illustrated by the weighted mean ɛHft of the sample population, and differences in the corresponding over-dispersion where the MSWD is in excess of that expected for a single population at the stated level of uncertainty. With an ɛHft uncertainty of ~ ± 0.5 ɛHft (2σ), the data show that the LHG and aplite dyke samples form reproducible single populations without over-dispersion which indicates that zircon crystallised from a melt with homogenous ɛHf, whereas the pre-mineralisation samples (> 167.4 Ma) show over dispersion that must result from variable ɛHf between zircon grains, which suggests isotopic heterogeneity within the melt. The mean value is consistent (within ~ 0.2 ɛHft) between syn-mineralisation samples yet is ~ 1 ɛHft lower in the oldest Bear QM sample. The increased range and lower ɛHft indicates greater crustal assimilation, which probably occurred during ascent into the sub-volcanic environment.
Given the paucity of any major component of zircon xenocrysts older than the Triassic volcanic and sedimentary country rocks (Figs. 4, 7; Supplementary Data 2 & 5), we infer that there was little continental crustal material present to impart large variations in ɛHft upon assimilation. Nevertheless, there is a systematic variation between earlier and later pre-mineralisation intrusions that is best explained by the pre-mineralising McLeod QMD and Bear QM magmas (prior to ~ 167.4 Ma) having undergone transport, storage and evolution within, and were contaminated by, the crustal column leading to the more varied and crustal isotopic signatures (Sr, O and Hf). In contrast, after ~ 167.4 Ma, these rocks no longer show this signature, indicating no discernible crustal assimilation during magma transport and storage. It suggests that the mineralising LHG-related melts reached their evolved compositions within a lower crustal environment where they were only exposed to homogenous, mantle-derived magmas. This supports different evolution zones for the pre- and syn-mineralisation melts. These could be either discretely located throughout the crust or within the same ‘hot-zone’67, reflecting melt extraction in variable proximity to the country rock, with LHG melts being entirely encapsulated by juvenile, mantle-derived rocks, with negligible assimilation of other crustal components prior to emplacement.
Depth of melt evolution
Since melt chemistry is partly controlled by the pressure of differentiation e.g.67, it can offer insights into the depth at which melts evolve. For example, increased Sr/Y, as observed in the LHG, is often used to infer a greater depth of fractionation due to an increased abundance of amphibole and suppressed plagioclase crystallisation within a relatively deep fractionating assemblage12,13,67,76. The compositions and normative mineralogy of H2O-saturated minima and eutectics for haplogranitic melts also share a relationship with pressure e.g.77,78. This pressure equates to the approximate depth at which the melt reached the eutectic, or evolved to its bulk composition, rather than the emplacement depth.
The normative mineralogy of the LHG, porphyry and aplite dykes show a close fit to the H2O-saturated minima and eutectics for haplogranitic melts e.g.77,78 (Fig. 9 & S18). LHG samples and porphyry dyke samples cluster between the ~ 450 MPa and 1000 MPa minima. Conversely, aplite dyke samples cluster between the ~ 75 MPa and 200 MPa minima. Assuming lithostatic pressure with an average overburden density of 2.5 g/cm3, determined pressures roughly equate to a melt evolution depth of ~ 20–40 km for the LHG and porphyry dykes, and ~ 3–8 km for aplite dykes.
As with the geochemical signatures (Figs. 5, 6, 7), which are indicative of a deeper amphibole-dominated evolution e.g.12,13, these melt minima relationships also support a deep melt evolution (~ 20–40 km) for the LHG, and also probably for the porphyry dykes (Fig. S18), rather than shallow fractionation at the emplacement level. However, this whole rock barometry indicates that aplite dykes (including those that contain textures indicating fluid exsolution and Cu mineralisation33,52; Fig. 3 & S3-S6), were sourced from shallow depths (~ 3–8 km), near their level of emplacement, probably representing late-stage melts associated with the LHG.
A rapid change in magmatic plumbing to tap porphyry ore-forming magmas
Previous models for the assembly of the Yerington batholith e.g.27,32,47 suggest sequential emplacement of the McLeod QMD and Bear QM, yet this is at odds with the revised chronology where these two intrusive units crystallised and must have been emplaced at least in part over the same ~ 800 kyrs time period (Fig. 4, 10). Whilst these units are mineralogically and texturally distinct27, their isotopic signatures and whole-rock and zircon geochemistry are very similar (Figs. 5, 6, 7, 8), suggesting similar sources and evolutionary pathways, likely in a mid-crustal storage zone (Fig. 10), although this must have evolved over the 1.5 Myrs of upper crustal activity. The whole-rock and zircon geochemical signatures of the pre-mineralisation McLeod QMD and Bear QM units (Fig. 5, 6, 7 & S9–S15) are consistent with clinopyroxene-plagioclase-dominated fractionation in the mid-crust (Fig. 10). This contrasts with the change to the syn-mineralisation signatures (Figs. 5, 6, 7 & S9–S15) which indicate an amphibole-dominated lower crustal evolution (~ 20–40 km depth; Figs. 9, 10). The intrusions either-side of this geochemical change are both relatively evolved and have comparable indicators of fractionation, such as whole-rock SiO2 and zircon Hf concentrations (Figs. 6, 7).
The shift in the compositions of the magmas which formed the pre-mineralising intrusions and then the LHG, along with the change in the dominant fractionating assemblage, is constrained to within < 200 kyrs and is coincident with the onset of porphyry mineralisation. For the LHG and porphyry dykes, although there is little difference in their whole-rock (Fig. 5 & S10), zircon trace element and isotopic compositions (Fig. 7, 8), and where they plot on the melt minima diagram (Fig. S18), there is no evidence from within the ~ 8 km depth of exposure that the porphyry dykes were derived from the upper parts of the LHG (as per previous models e.g.27). Instead, the porphyry dykes may reflect the same or similar intrusive events that formed the LHG. Once emplaced at shallow crustal levels, the LHG magmas underwent further fractionation (at ~ 3–8 km depth, based on melt minima plots; Fig. 9), potentially forming igneous banding textures (Fig. S2), to form more evolved and volatile-rich melts that were episodically injected as aplite dykes over a period of at least ~ 400 kyrs (Fig. 4). Multiple, episodically emplaced generations of these aplite dykes, which provide textural evidence for undercooling and exsolution of mineralising fluids (mineralised miarolitic cavities33,52), are associated with early A-type veins and likely acted as crystal mush conduits for mineralising fluids33 (Fig. 3c–f & S3–S6). Zircon U–Pb and molybdenite Re-Os ages indicate hydrothermal mineralisation occurred episodically over time-periods potentially in excess of 1.5 Myrs post-emplacement of the LHG cupola (Fig. 4), perhaps coincident with the eruption of the younger propylitically-altered components of the Fulstone Volcanics that bear the same zircon geochemical signatures as the LHG (Figs. 4, 7, S7, S12-S15). The proposed time-period for porphyry ore formation, which may have exceeded 1.5 Myrs post emplacement of the LHG cupola, is not uncommon for medium to large scale, composite porphyry systems e.g.9,79,80,81,82, or where multiple, episodic ore-forming magmatic-hydrothermal events are documented e.g.72,81,83. This all suggests that relatively evolved, internal or deep parts of the LHG remained active and continued to produce the magmas and associated magmatic-hydrothermal fluids responsible for porphyry deposit formation after emplacement and crystallisation of the cupolas and upper regions of the LHG (Fig. 10). It also implies that the LHG pluton was episodically recharged rather than being emplaced as a single intrusive event, as previously suggested27,38.
Given the similarities in zircon trace element geochemistry between the mineralised porphyry dykes from the Ann Mason and Yerington porphyry deposits (Figs. S13–S15;36), and their mineralogy27, they are probably genetically related. By extrapolation, this is also likely to be the case for porphyry dykes in the Yerington Districts’ two other known porphyry deposits: Bear and MacArthur. Because of this, it is probably salient for future computational simulations of batholith construction and mineralisation to include fluids derived from across all porphyry centres; this will yield a considerably larger copper endowment than when individual porphyry centres are considered (> 9 Mt of contained Cu42,43,44,45).
The abrupt (~ 100–200 kyrs) change to geochemical signatures indicative of magmas from a lower crustal amphibole-stable, plagioclase-supressed, evolution zone (from whole-rock Sr/Y and REE patterns, and zircon geochemistry, Figs. 5, 6, 7), in tandem with an increase in ore-forming potential, requires an explanation. It is plausible that these changes could have occurred in a single magma reservoir as a result of a progressive long-term transition. Within the lower crust, the rapid change in magma chemistry could reflect a relatively discrete temporal point at which the ‘amphibole-in’ line was suddenly crossed. This could occur either due to a build-up of volatiles following fractionation of anhydrous phases, over a period of at least 1.5 Myrs (Fig. 4), or because of an injection of new melts into a lower crustal clinopyroxene cumulate pile or ‘sponge’ that reacts with new melt to become progressively replaced by amphibole76. The dated porphyry dyke that sits at the temporal onset of mineralisation (AC25; Fig. 4), and has zircon geochemistry appearing to ‘straddle’ the pre-and syn-mineralisation signatures, could mark this threshold being crossed in a transitional phase of magmatism, although this is a feature common to all porphyry dykes, regardless of their timing (Fig. 7 & S13–S15).
Although we cannot rule out a model where the change captures a single petrological event in a transitional process, there are several features that do not support progression within a single magma evolution zone. If the change were merely due to a transition in the magma supplied to the upper crust we would perhaps not expect the sharp contacts between the mineralogically distinct McLeod QMD and LHG plutons27. A scenario where the fractionating assemblage suddenly changes in a single transitional melt extraction zone is also challenging to reconcile given the variations observed in the isotopic data (Fig. 8), i.e. from a heterogeneous distribution, indicative of variable interaction with crustal components, to a homogeneous, less contaminated, mantle-derived signature within < 200 kyrs. Instead, magmas being sourced from discrete melt evolution zones within the crust, with pre-mineralisation intrusions evolving at shallower levels, surrounded by country rocks, and the ore-related intrusions evolving within a deeper zone dominated by mantle-derived rocks would be a better fit to the data (Fig. 10). This idea is also supported by the geochemical indicators of melt evolution depth (Figs. 5, 6, 7, 9 & S18). We envisage that the earlier, pre-mineralisation stage magmas were derived from the mid-crust. During protracted storage and evolution, these assimilated crustal materials. The magmatic plumbing then shifted to tap magmas from a deeper, lower crustal ‘hot-zone’ (~ 20–40 km; Fig. 9, 10)67, which likely evolved over extended time periods. In this scenario, the pre-mineralisation geochemical signature of the zircon cargo of the porphyry dykes would be acquired as they intruded up through the pre-cursor magmatic system on route to their level of emplacement. It is also feasible that the mid-crustal melt evolution zones of the McLeod QMD and Bear QM could have remained active post emplacement, or during the evolution of the syn-mineralisation magmas. In addition, the geochemical differences do not exclude progressively more oxidising conditions within the magmatic system e.g.29, but this is unlikely to have controlled all the observed changes.
Genetic implications for porphyry deposit-forming magmatic systems
The apparent change in geochemistry (whole-rock and zircon; Figs. 5, 6, 7, 8) as the Yerington system began to produce porphyry deposits is consistent with observations from a wide range of similar magmatic centres globally where precursor magmatism and syn-mineralisation intrusions have been examined e.g.19,22,24,26,29,84. Typically, these changes have been interpreted solely as being due to long-term, tectonically driven arc-scale, transitional processes over millions of years or ‘ramp-ups’ towards ore-formation e.g.11,16,21,23. However, these explanations are relatively poorly constrained due to limited exposure in most porphyry systems30. From our studies of the well exposed, ~ 8 km deep cross section through the Yerington system, the possibility exists for a much more rapid (< 200 kyrs) shift to porphyry deposit-forming magmatism. Significant changes in geochemical signatures over relatively short timescales at the transition to porphyry deposit forming magmatism have been indicated elsewhere e.g.24,84, but this study provides unprecedented temporal and spatial resolution due to the acquisition of our new high precision geochronological framework and the unique depth constraints at Yerington. This short timeframe does not necessarily contradict the suggestion of longer-term progressions towards ore-forming arc magmas, commonly seen in other systems. Rather it captures how rapid changes in the ore-forming potential of the magmatic system may occur. In other porphyry centres, the rapid timescales are often inconspicuous due to the limited rock record available. As such, the much longer durations between precursor and ore-related magmatism documented elsewhere, alongside their corresponding change in geochemistry, may relate to differences in the juxtaposition of upper crustal magmatic expressions over the protracted duration of the magmatic system. For example, when only the shallow levels (e.g. ~ 2 km palaeodepth) are considered at Yerington, porphyry dykes that share a comparable magmatic evolution to the LHG could yield an apparent temporal difference of ~ 1.7 Myrs with the Bear QM they cross-cut. The timescale of the change in geochemical signatures between these units will appear drawn-out and to have developed over longer timescales, whereas at depth the system is demonstrably more concurrent.
The recognition of rapid changes within the magmatic plumbing system requires a new perspective when interpreting magmatic processes in ore-forming systems. Because the magmas responsible for ore-formation underwent different routes of evolution and likely were tapped from spatially independent, deeper melt zones, it suggests the processes and evolution histories of early-intruded plutons cannot necessarily be used to infer whether other parts of the batholith may have produced porphyry-type deposits, and we advise caution over the use of earlier parts of the magmatic system to infer the nature of what has been removed or added to the melts over longer time periods or apparent progressions of melt chemistry such as metal contents that may be removed by earlier sulphide fractionation e.g.85. With lower precision geochronology (e.g. 2% typical of microbeam U–Pb methods), these earlier intrusive phases that emanate from potentially disparate magmatic plumbing systems may even appear to be ‘coeval’ with mineralisation.
The short, < 200 kyrs timescale for the emergence of the geochemical signatures associated with mineralisation that appeared throughout the magmatic system (in plutons, dykes and volcanics) significantly narrows and better defines the temporal footprint that can be used to identify ore-forming processes within the rock record. This has significance in the development and refinement of porphyry exploration indicators by increasing the potential spatio-temporal efficacy of using these geochemical ‘fertility’ signatures to isolate areas most prospective for porphyry-style mineralisation. Whilst the large-scale long-duration, tectonically driven signatures previously identified can still be critical in defining general targets, increased resolution by which the ore forming signature can be discriminated can lead to greater confidence in identifying and discovering the next generation of porphyry deposits, which are likely to be deeper and often under cover and so will be more difficult to find10.