Pressure, temperature and fluid conditions during emerald

precipitation, southeastern Yukon, Canada: fluid inclusion and

stable isotope evidence

Dan Marshalla,*, Lee Groatb, Gaston Giulianic, Don Murphyd, Dave Matteye,T. Scott Ercitf, Michael A. Wiseg, William Wengzynowskih, W. Douglas Eatonh

aEarth Sciences Department, Simon Fraser University, Burnaby, BC, Canada V5A 1S6bEarth and Ocean Sciences, University of British Columbia, Vancouver, BC, Canada V6T 1Z4

cCentre de Recherches Petrographiques et Geochimiques, UPR A 6821, 15 rue Notre Dame des Pauvres, BP 20,

54501 Vandoeuvre-les-Nancy Cedex, FrancedYukon Geology Program, Government of the Yukon, Whitehorse, YT, Canada Y1A 2C6

eDepartment of Geology, Royal Holloway, University of London, Egham Hill, Egham, Surrey TW20 OEX, UKfResearch Division, Canadian Museum of Nature, Box 3443, Station D, Ottawa, ON, Canada K1P 6P4

gSmithsonian Institute, Mineral Sciences NHB119, Washington, DC 20560-0119, USAhExpatriate Resources Limited and Archer Cathro & Associates (1981) Limited, 1016-510 West Hastings Street,

Vancouver, BC, Canada V6B 1L8

Received 30 November 2001; received in revised form 3 May 2002

Abstract

The Crown emerald veins are somewhat enigmatic, displaying characteristics that are common to emerald deposits of

tectonic–hydrothermal origin and of igneous origin. The veins cut the Fire Lake mafic meta-volcanic rocks, occurring within

600 m of an outcrop of Cretaceous S-type granite. Field work and vein petrography are consistent with a polythermal origin for

the veins. The primary vein mineralogy is quartz and tourmaline with variable sized alteration haloes consisting of tourmaline,

quartz, muscovite, chlorite and emerald. The veins weather a buff brown colour due to jarosite, scheelite and minor

lepidocrocite, which were precipitated during the waning stages of vein formation. Microthermometic studies of primary fluid

inclusions within emerald growth zones are consistent with emerald precipitation from H2O–CO2–CH4 (FN2FH2S) bearing

saline brines. The estimated fluid composition is approximately 0.9391 mol% H2O, 0.0473 mol% CO2, 0.0077 mol% CH4 and

0.0059 mol% NaCl (f 2 wt.% NaCl eq.). Fluid inclusion and stable isotope studies are consistent with vein formation in the

temperature range 365–498 jC, with corresponding pressures along fluid inclusion isochore paths ranging from 700 to 2250

bars. These data correlate with a very slow uplift rate for the region of 0.02–0.07 mm/year.

Emerald deposits are generally formed when geological conditions bring together Cr (FV) and Be. Cr and V are presumed

to have been derived locally from the mafic and ultramafic rocks during hydrothermal alteration. The Be is most likely derived

from the nearby Cretaceous granite intrusion.

D 2002 Elsevier Science B.V. All rights reserved.

Keywords: Fluid inclusions; Stable isotopes; Emerald; Thermobarometry; Yukon

0009-2541/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved.

PII: S0009 -2541 (02 )00277 -2

* Corresponding author. Fax: +1-604-291-4198.

E-mail address: marshall@sfu.ca (D. Marshall).

www.elsevier.com/locate/chemgeo

Chemical Geology 194 (2003) 187–199

1. Introduction

During the summer of 1998, emerald was discov-

ered at the Crown showing of the Goal Net property, 7

km north of Fire Lake in the southeastern Yukon (Fig.

1). Although there are numerous occurrences of beryl

within the Canadian Cordillera, there are only three

documented emerald showings. These are the Crown

(Yukon), Lened (Northwest Territories) and Red

Mountain (near Stewart, British Columbia). The latter

two have V-bearing emerald with the Goal Net prop-

erty being the only occurrence in the Canadian Cor-

dillera of traditional Cr-bearing emerald.

Emerald is thought to be rare because Be and Cr

are not commonly found together in sufficient con-

centrations in iron poor medium temperature environ-

ments to stabilize Cr–emerald. Mafic and ultramafic

rocks are generally enriched in Cr and V, whereas

rock-enriched Be, Al and Si are pegmatites, evolved

granites, metamorphic rocks and black shales. Exten-

sive work (Schwarz and Giuliani, 2001; Banks et al.,

2000; Giuliani et al., 2000; Vapnik and Moroz, 2000)

has shown that there are two general models for

emerald precipitation: magmatic associated and tec-

tonic–hydrothermal.

Emerald deposits linked to magmatism are gen-

erally associated with granitic intrusions and host

rocks ranging from sedimentary to volcanic rocks.

The source of Be for the emerald is generally

surmised to be the granitic rocks with some con-

tribution from sedimentary country rocks. In this

model, Cr is exclusively derived from the country

rocks, which are generally mafics, ultramafics and

shaley sediments. Some examples of magmatic-asso-

ciated emerald are Ndola (Zambia), Belingwe (Zim-

babwe), Mananjary (Madagascar), Carniba-Socoto

(Brazil), Khaltaro (Pakistan) and Hiddenite (USA).

The emerald occurs in veins and associated alter-

ation haloes, and are generally precipitated at tem-

peratures in excess of 350 jC from H2O–CO2FCH4 saline brines with salinities varying from 12 to

40 wt.% NaCl equivalent (wt.% NaCl eq.). The

emerald is precipitated during metasomatic alteration

of the host rock by high-temperature fluids derived

from the cooling intrusions (Schwarz and Giuliani,

2001).

The emerald associated with tectonic–hydrother-

mal activity is generally located along major crustal

faults or shear zones in mafic to ultramafic schists.

These structures localize fluid flow with emerald

precipitating at the alteration fronts of rock–fluid

interaction zones. Generally, the metamorphic grade

of the host rocks varies from upper greeenschist to

lower amphibolite facies. In this type of deposit, the

Cr is derived locally with emerald precipitating when

a Be-rich fluid encounters the Cr-bearing schist. The

fluids responsible for tectonic–hydrothermal emerald

deposits, in general, range from low salinity (Moroz

and Vapnik, 1999, 2001) to supersaturated saline

fluids (Banks et al., 2000; Schwarz and Giuliani,

2001). Sulphate reduction of organic matter in shales

is also thought to play an important role in buffering

fluid chemistry during the precipitation of emerald

(Ottaway et al., 1994; Giuliani et al., 2000).

2. Regional geology

The emerald mineralization at the Crown showing,

of the Goal Net property, is hosted within the alter-

ation haloes of quartz–tourmaline veins that crosscut

the Devono-Mississippian Fire Lake metavolcanic

rocks (Fig. 1). This is a diverse unit comprised of

boninites, low-Ti basalts, normal MORB basalts, and

LREE-enriched tholeiites (Piercey et al., 1999). Basal

members within this unit are generally meta-pelites

with localized calcareous beds (Murphy and Piercey,

2000; Murphy, 1998). The host rocks to the emerald

are meta-boninites. Intruding into the metavolcanic

rocks are partially metamorphosed mafic and ultra-

mafic rocks. All these rocks have been metamor-

phosed to greenschist facies and have a well-

developed planar fabric dipping slightly northward.

Overlying the Fire Lake rocks are the Mississippian

Grass Lakes succession consisting of feldspar–mus-

covite–quartz schist of volcanogenic and volcaniclas-

tic origin and a grey calcareous phyllite unit. This

entire package of rocks was deformed and metamor-

phosed during late Paleozoic. Small mafic dykes were

emplaced during the Triassic, with further deforma-

tion, metamorphism and imbrication during the Creta-

ceous. A series of Cretaceous two-mica (S-type)

granitic intrusions mark the waning stages of defor-

mation and metamorphism at 112 Ma (U–Pb, zircon,

Mortensen, 1999). This entire package (Yukon–

Tanana terrane) is cut 14 km to the southwest by the

D. Marshall et al. / Chemical Geology 194 (2003) 187–199188

Fig. 1. Location map of the emerald veins showing the major rock types and geological relationships within the study area.

D.Marsh

allet

al./Chem

icalGeology194(2003)187–199

189

Tintina fault system, a major crustal feature with 450

km of right lateral displacement (Roddick, 1967).

There are eight distinct emerald zones within a few

hundred square meters. The closest Cretaceous granite

outcrop occurs 600 m to the southeast of the emerald

zones, but field mapping suggests the Cretaceous

granite may be present approximately 500 m below

the emerald veins. There are numerous generations of

quartz veins in the general area of the granite and the

emerald mineralization. These range from pure quartz

veins of 1 m in width to narrower quartz–tourmaline

veins. Due to the limited exposure, the maximum

length of the quartz veins has not been determined.

Some quartz and quartz–tourmaline veins cut the

granite indicating that at least some of the veins

postdate the granite. No veins hosting emerald have

been traced into the granite and there is no field

evidence suggesting a relationship between the emer-

ald and the granite. On the other hand, none of the

geological evidence suggests that the emerald miner-

alization is not related to the granite.

3. Petrology

The emerald zones consist of quartz–tourmaline

veins with a muscovite–tourmaline alteration halo

(Fig. 2). The mineralogy within the halo consists of

chlorite–muscovite– tourmaline–quartz–scheelite–

jarosite. The emerald occurs in vug-like structures

almost exclusively within the alteration halo and

rarely in the quartz vein. The alteration halo is up to

1.5 m thick but may be absent in places. The emeralds

are up to 4 cm in length. They are a deep green colour

and have good clarity.

Petrographic examination of the veins suggests that

the primary mineral assemblage within the veins is

quartz–tourmaline. At the vein–halo interface, the

stable mineral assemblage is quartz– tourmaline–

muscovite–chlorite–emerald. Jarosite, lepidocrocite

and scheelite are present in the alteration halo, but

appear to be late in the alteration assemblage. Schee-

lite fills cracks in brittlely deformed tourmaline and is

interpreted to have precipitated during the waning

stages of hydrothermal alteration. The jarosite is fine

grained, and when altered, gives the haloes a yellow-

ish colour and a clay-like appearance where present in

significant quantity.

The emerald has been studied in detail by Groat et

al. (in press). It is Cr-bearing with concentrations of

up to 7816 ppm and V contents in up to 333 ppm.

These values are similar to a number of emerald

deposits worldwide, most notably those of Pakistan

and Somondoco, Colombia. The tourmaline proximal

to the veins is generally dravite, while the distal

tourmaline found in quartz veins and in the granite

is usually schorl or dravite. Tourmaline found as solid

inclusions within the emerald has compositions, deter-

mined by electron microprobe analyses, ranging from

dravite to uvite. Other solid inclusions found within

Fig. 2. Schematic cross-section of a quartz tourmaline vein showing the relationship to the overlying ultramafic rocks, alteration halo and

emerald distribution. Qtz: quartz, Tm: tourmaline, Em: Emerald, Chl: Chlorite, Sch: Schist.

D. Marshall et al. / Chemical Geology 194 (2003) 187–199190

the emerald, using scanning electron microscopy, are

calcite, chalcopyrite, molybdenite, phlogopite, pyrite,

quartz, scheelite and zircon.

Fluid inclusions are abundant within the emerald,

and in some rare cases, define growth zones (Fig.

3). The bulk of the fluid inclusions within the

emerald occurs along healed fractures or exist as

isolated inclusions. The inclusions generally display

consistent phase ratios (Fig. 4), with some clear

evidence of post-entrapment volume changes in

some inclusions. Despite the presence of abundant

solid inclusions within the host emerald, there have

been very few accidental (and no daughter) inclu-

sions identified. The accidental solids are generally

a transparent birefringent micaceous mineral. The

dominant fluid phase is an aqueous brine occupying

approximately 65% of the fluid inclusion volume.

The other two fluid phases are liquid and gaseous

carbonic fluids, each representing approximately 28

and 7 vol.%, respectively. Fluid inclusions hosted

within quartz display variable phase ratios. In most

cases, clear evidence of necking down can be seen.

This has resulted in inclusions consisting of high

proportions of carbonic fluid and corresponding

fluid inclusions with high H2O/CO2. No evidence

of a boiling fluid was observed and the rather

constant H2O/CO2 in the emerald is consistent with

a fluid trapped in the one-phase field.

Two types of fluid inclusions have been observed,

hosted within the quartz of the nearby Cretaceous

granite. These small (f 10 Am) melt inclusions

represent primary fluid inclusions within the quartz,

as they were clearly trapped as the quartz crystallized

in equilibrium with a melt. These inclusions have a

relatively small (f 3 vol.%) shrinkage bubble. This is

indicative of a melt that has low quantities of volatile

gases (Roedder, 1984). A second type of fluid inclu-

sion exists within the quartz from the granite, occur-

ring as secondary fluid inclusions along healed

fractures. These inclusions contain a liquid and a

Fig. 3. Photomicrograph showing textural equilibrium between quartz (Qtz), tourmaline (Tm) and emerald (Em). Growth zones (GZ) within the

emerald are rare and are composed of alternating bands of fluid-inclusion-free and fluid-inclusion-rich emerald. Photo taken under cross nicols.

D. Marshall et al. / Chemical Geology 194 (2003) 187–199 191

vapour phase at room temperature. The vapour occu-

pies approximately 5 vol.% of the inclusion at room

temperature. These two-phase inclusions have not

been observed in the emerald veins.

4. Emerald vein fluid inclusion data

The fluid inclusion data from the emerald veins

are derived from inclusions hosted within emerald.

Reconnaissance observations of three-phase fluid

inclusions hosted within quartz from the emerald

veins indicate that the quartz hosted fluid inclusions

contain the same fluid as the inclusions hosted

within the emerald. Due to the petrographic evi-

dence of post-entrapment volume changes within

the quartz-hosted fluid inclusions, it was deemed

superfluous to collect data from the quartz. Hence,

the data and interpretation is based upon the fluid

inclusions in emerald. However, this data can be

Fig. 4. Typical three-phase fluid inclusions hosted within emerald from three different emerald (Em) specimens. (A) Sample 10-1-C1: slightly

stretched fluid inclusion after heating to approximately 350 jC. (B) Relatively flat fluid inclusion, used for initial estimate of vapour fraction

within the fluid inclusion population, showing the relative proportions of Aqueous Liquid (Laq), Carbonic Liquid (Lc) and Carbonic Vapour

(Vc). (C) Sample 10-1-C2: three phase inclusions prior to any measurements showing consistent phase ratios between different fluid inclusions.

All photos taken at room temperature in plane polarized light.

D. Marshall et al. / Chemical Geology 194 (2003) 187–199192

extrapolated to the quartz as petrographic and fluid

inclusion evidence indicates that quartz coprecipi-

tated with the emerald.

Microthermometric data were obtained from dou-

bly polished mineral plates utilizing a Linkam THMS-

G 600 heating–freezing stage on an Olympus BX50

microscope equipped with 40� and 100� long

working distance Olympus objectives. The stage is

capable of measurements in the � 190 to + 500 jCrange. The stage was calibrated with two synthetic

fluid inclusions having four readily observable phase

transitions at � 56.6, 0.0, 10.0 and 374.1 jC. Thestage was periodically checked against the standards

and results were always within F 0.1 jC of the two

low temperature standards and within 0.5 jC of the

higher temperature phase transition. From these meas-

urements, we conclude that the (2r) error is less thanF 0.1 jC for temperatures below 50 jC and 0.5 jCfor temperatures in excess of 50 jC.

Upon cooling from room temperature to � 190 jC,the three-phase fluid inclusions nucleate three addi-

tional phases. At approximately � 36 jC, clathratenucleates. This is followed closely by the formation of

ice at approximately � 48 jC. Solid CO2 nucleates at

approximately � 101 jC.Heating of the fluid inclusions from � 190 jC

results in numerous phase changes. The first occurs at

approximately � 70 jC. The solids in the inclusions

darken slightly, and occasionally, there are cracks

developed in the solids. Solid CO2 melts over the

temperature range � 64 to � 56.6 jC (Fig. 5, Table

1). Continued heating results in continuous melting of

ice in the inclusions commencing at an average

eutectic temperature of � 20.1 jC, with ice melting

temperatures occurring over the temperature interval

� 8.1 to � 0.5 jC. Further heating reveals final

clathrate melting temperatures in the 7.4–12.1 jCrange (Fig. 6). More heating results in the homoge-

nization of the carbonic fluids into the vapour (dew-

point transition) over the temperature range 22.3–

29.8 jC. Final homogenization to the liquid takes

place over the temperature range 259–367 jC (Fig.

7). Many of the fluid inclusions decrepitate prior to

total homogenization, and some inclusions show evi-

dence of stretching such as cracks and enlarged

vapour bubbles. Smaller inclusions are less likely to

stretch and our observations are consistent with total

homogenization temperatures in the range 260–310

jC, with minor stretching generally accounting for the

higher homogenization temperatures.

Freezing point depression of solid CO2 within the

inclusions is indicative of additional gas species. The

two most likely gases are CH4 and N2. Clathrate

melting temperatures in excess of 10.0 jC are indi-

cative of CH4, as its presence expands the clathrate

stability field (Diamond, 1992). N2 has the opposite

effect lowering the clathrate melting temperature. The

high clathrate melting temperatures are indicative of

CH4 in the inclusions, but do not negate the presence

of N2. Raman studies to determine the gas composi-

tions of the fluid inclusions in emerald have been

unsuccessful as the emerald fluoresce and mask any

Raman peaks. Crushing studies in paraffin and glyc-

erol have been inconclusive in determining gas com-

positions.

It is evident from the variable solid CO2 and

clathrate melting temperatures that salinity and X-

CH4 vary within the fluid inclusion population. The

presence of methane, and thus elevated clathrate

melting temperatures, makes salinity determinations

difficult. The variable solid CO2 melting temperatures

(Table 1) are consistent with a range of X-CH4 in the

carbonic phase ranging from 0.0 to over 50 mol%

Fig. 5. Histogram showing the range of solid CO2 melting

temperatures for fluid inclusions hosted in emerald. n=Number of

inclusions measured.

D. Marshall et al. / Chemical Geology 194 (2003) 187–199 193

(Burruss, 1981). There are three inclusions (Table 1)

that have the theoretical solid CO2 melting temper-

ature (� 56.6 jC) in the H2O–CO2 binary system,

and these can be used to establish a minimum salinity.

One inclusion (10-1-C1-23) has a clathrate melting

temperature of 10.0 jC, which is consistent with a

pure H2O–CO2 fluid, i.e. 0 wt.% NaCl eq. Likewise,

ignoring the possible contribution of nitrogen and

acknowledging the fact that suitable experimental data

does not exist for our system; we can determine an

‘‘estimated’’ maximum salinity from the clathrate

melting temperature depression (due to the presence

of NaCl). To do this, we assume that the fluid

inclusion (10-1-C1-4, Table 1) with the highest clath-

rate melting temperature (12.1 jC) represents the

lowest salinity methane-bearing fluid inclusion. Fluid

inclusion 9-2 (Table 1) has the lowest clathrate melt-

ing temperature (7.4 jC), and thus maximum salinity.

Our ‘‘estimate’’ also assumes that clathrate melting

depression, due to salts, can be approximated by the

Table 1

Microthermometry of fluid inclusions hosted in emerald

Sample number Tm-CO2Te-ice Tm-ice Tm-clathrate Th-CO2

Th-tot

10-1-C1-1 � 59.1 � 2 9.7 27.6 287

2 � 59 9.7 26.6 310.2

3 � 59.1 � 2.8 9.2 28.1 Td>290

4 � 59.1 � 4.8 12.1 365.0

5 (necking) � 58 10.5 23.8 297.3

6 � 57 11.2 25.4 295.1

7 � 60 10.5 300

8 � 60 � 3.5 11.2 Td>290

9 � 57.4 � 3.2 9.9 27.8 290.5

10 � 57 � 0.5 9.9 27.7 291.8

11 � 57.9 9.8 27.5 >365

12 � 57.7 9.6 27 >365

13 � 59.3 10.6 Td>290

14 � 2.3 9.5 28.1 308.2

15 � 1.9 9.9 29.8 Td>290

16 � 58.1 � 3 10.7 23.7 294.6

17 � 59.7 � 5 10.9 259

18 � 58.4 � 4.5 11.5 353.6

19 � 56.1 � 3.7 10.5 28 310.8

20 � 56.6 � 4.1 10.2 310.1

21 � 57.1 � 4.4 10 22.3 341

22 � 58.9 � 4.1 11 >365

23 � 56.6 � 5.2 10 26.4 313.8

24 � 4.4 9.3 26.5 324.2

25 � 56.6 � 4.6 9.7 25.3 324.9

26 � 4.4 9.9 27.8 >365

27 � 35 � 4.4 9 27.5 301.5

28 � 17 � 4.4 10 290.9

30 � 57.2 � 4.3 8.9 27.4 Td>300

31 � 58.5 � 4.7 11.1 366.9

32 � 56.7 � 3.5 10.1 367

33 � 58.2 353

10-1-C2-1 � 57.1 � 22 � 4 10.1 24.8 288

10-1-C2-2 � 57.3 � 17 � 1.7 10.8

10-1-C2-3 � 58.1 � 1.7 10.5

10-1-C2-4 � 57 � 4 10.1 24 294

9-1 � 59.3 � 17 � 7 8.5 25.7

9-2 � 59.5 � 21 � 8 7.4 28

9-3 � 63 � 23 � 7.9 8.7

9-4 � 62.6 � 17 � 8.1 8.1 28.3

9-5 � 62 � 18 9.4 24.2

9-6 � 62 � 15 � 8 9.1 24

9-7 � 63.5 � 17 � 7.4 8.7 25.7

9-8 � 27 � 7 10

9-9 � 63.1 � 5.2 8.7

9-10 � 61.7 9.6

9-11 � 61.6 � 19 � 6.9 8.6

9-12 � 61.3 9.8

9-13 � 62.4 9.8

9-14 � 61.2 � 6.8 9.1 28.3

9-15 � 60.4 � 31 � 6 8.6 29.4

Averages � 59.3 � 20.1 � 4.2 9.8 26.6 313.5

Standard deviation 2.4 18.4 2.0 0.9 1.8 28.9

Sample number Tm-CO2Te-ice Tm-ice Tm-clathrate Th-CO2

Th-tot

Minimum � 63.5 � 35.0 � 8.1 7.4 22.3 259

Maximum � 56.6 � 15.0 � 0.5 12.1 29.8 367

Table 1 (continued )

Fig. 6. Histogram of clathrate melting temperatures. Temperatures

in excess 10 jC are generally attributed to the melting of CH4–

calthrates. n=Number of measurements.

D. Marshall et al. / Chemical Geology 194 (2003) 187–199194

pseudo-ternary (H2O–CO2–NaCl) system with an X-

CH4 of 0.1 (Diamond, 1992) in the carbonic compo-

nent. Using the data of Diamond (1992), a clathrate

melting temperature of 7.4 jC corresponds to an

‘‘estimated’’ maximum salinity of approximately 8

and an average salinity 2 wt.% NaCl eq. Additionally,

there is a discrepancy between our Th-CO2and the

theoretical maximum of 23.2 jC for the pseudo

ternary H2O–CO2 with an X-CH4 of 0.1 (Diamond,

1992). The increase of the Th-CO2beyond the critical

point of the pseudo-binary can be accounted for by

minor amounts of H2S. Diamond’s (1992) data was

extrapolated beyond the critical curve for the salinity

estimates for our inclusions.

Pending a Raman analyses of the carbonic phase,

vapour fractions and initial X-CO2 within the fluid

inclusions were estimated visually and then refined

using the technique of Diamond (2001) and Bakker

and Diamond (2000), assuming all compressible gases

within the inclusions behave as CO2. Finally, X-CH4

was determined from the average depression of the

CO2 triple point and the initial X-CO2 was subse-

quently corrected to account for the methane with the

inclusions, using the method of Diamond (2001). X-

CO2 in the aqueous phase is estimated at 0.027. This

is estimated from the maximum CO2 solubility in the

aqueous phase in the H2O–CO2 binary. This is a

minimum, as a higher internal (fluid inclusion) pres-

sure due to methane will enhance the solubility of

CO2. X-CH4 in the aqueous phase is assumed to be 0.

Internal fluid inclusion pressure at the average final

clathrate melting temperature is estimated at 46 bars

from Diamond (1992). Using these assumptions,

approximations and the data obtained from micro-

thermometry (Table 1), an ‘‘estimated’’ average fluid

inclusion composition of 0.9391 mol% H2O, 0.0473

mol% CO2, 0.0077 mol% CH4 and 0.0059 mol%

NaCl (f 2 wt.% NaCl eq.) and an estimated bulk

molar volume of 26.35 cm3/mol was determined.

To complement the microthermometric salinity

determinations, three grains of emerald and quartz

were frozen in liquid nitrogen, broken and rapidly

placed (uncoated) in the chamber of a Bausch and

Lomb Nanolab scanning electron microscope (SEM)

equipped with an energy dispsersive X-ray (EDX)

system. The mineral grains were examined using a

modified technique after Kelly and Burgio (1983) for

breached fluid inclusions. Several cavities resembling

breached fluid inclusions were found in the quartz and

emerald samples. No residua was visible within or

surrounding the fluid inclusion cavities. The electron

beam was directed inside a number of breached

inclusions and an EDX spectrum collected. Likewise,

an area encompassing the breached inclusions and the

surrounding area was scanned and an EDX spectrum

collected. In all cases, the spectra contained the

theoretical peaks of the host minerals, but no peaks

attributable to the major chlorides (Na, K, Mg, Ca or

Cl) were observed.

5. Cretaceous granite fluid inclusion data

Attempts to crush the rare primary melt inclusions

within quartz from the granite failed to identify any

compressed gases within the shrinkage bubbles. The

relatively small volume of the shrinkage bubble

makes this test difficult; however, repeated tests

yielded consistent data. Reconnaissance microther-

mometry of the secondary fluid inclusions indicates

Fig. 7. Histogram of total homogenization temperatures. All

inclusion homogenized to the vapour phase. The secondary peak

at higher temperatures is attributed to a volume change within the

fluid inclusion due to the increased internal pressure at high

temperature. Some inclusions underwent total decrepitation (Td) at

temperature in excess of 300 jC.

D. Marshall et al. / Chemical Geology 194 (2003) 187–199 195

they contain no compressible gases and undergo total

homogenization approximately 50j less than the

emerald inclusions, suggesting that these fluids are

markedly different from the fluid inclusions associ-

ated with the emerald veins.

6. Stable isotopes

Oxygen isotope data were obtained from scheelite,

quartz and tourmaline separates from a single-hand

specimen in an attempt to establish if these minerals

were contemporaneous and to constrain the temper-

atures of vein formation. Petrographic relations were

consistent with the coprecipitation of quartz and

tourmaline, followed by scheelite. The analytical

conditions are described in Mattey et al. (1994) and

Mattey and MacPherson (1993). The isotope data are

presented in Table 2. Quartz–tourmaline thermometry

based on the calibration of Kotzer et al. (1993) yields

equilibrium temperatures ranging from 365 to 498 jC.The quartz–scheelite fractionation curve of Zheng

(1992) yields temperatures of 248–263 jC. Tourma-

line–scheelite pairs yield temperatures in the range of

120–250 jC using the combined calibrations of

Zheng (1992, 1993).

Groat et al. (in press) present oxygen isotope data

for emerald from the Crown showing. Combining

their emerald data with the quartz isotope data from

this study yields variable equilibrium temperatures in

the range 569–1670 jC using the beryl–H2O cali-

bration of Taylor et al. (1992) combined with the

quartz–H2O calibration of Clayton et al. (1972).

These temperatures are discordant with temperatures

obtained for quartz–tourmaline pairs and inconsistent

with the metamorphic grade observed in the rocks.

This is good evidence for localized oxygen isotope

equilibrium throughout the deposit as the quartz–

tourmaline and beryl (emerald) samples were taken

from different localities within the showing.

7. Pressure–temperature conditions of vein

formation

The main purpose of the fluid inclusion study on

the emerald was to determine the composition of the

fluids responsible for the growth of emerald, to

determine isochoric constraints for these fluids and

to integrate these findings with the preliminary oxy-

gen–isotope thermometry to approximate the pres-

sures and temperatures of vein formation and dis-

tinguish between a tectonic–hydrothermal versus a

magmatic origin.

The fluid inclusion composition is rather complex

with an aqueous brine phase and a carbonic phase

containing CO2 and CH4. It may also contain minor

amounts of H2S and N2. In an attempt to generate

fluid inclusion isochores for an estimated average

composition, a variety of methods have been em-

ployed. Initially, the program FLINCOR (Brown,

1989) was used to calculate isochores. The Bowers

and Helgeson (1983) equation of state for H2O–

CO2–NaCl fluids was used, assuming all the carbonic

component of the inclusions was CO2. Secondly, an

isochore was calculated using the Jacobs and Kerrick

(1981) equation of state for H2O–CO2 fluids. A third

isochore was calculated using the GASWET8 pro-

gram and the Bakker (1999) equation of state using

the composition as determined in the H2O–CO2–

CH4–NaCl ternary (Fig. 8). Lastly, interpolation of

the experimental data of Gehrig (1980) was used to

derive an isochore in the H2O–CO2–NaCl ternary.

Repeated calculations with the Bakker (1999) equa-

tion revealed that it provides the best fit to the

experimental data of Gehrig (1980) for compositions

and molar volumes similar to the fluid inclusions in

this study. Therefore, our preferred isochoric path

would lie in the area bounded by the Gehrig inter-

polation and the Bakker (1999) equation of state

(Fig. 8).

Pressure–temperature (P–T) constraints can be

derived by intersecting the fluid inclusion isochores

with the temperature constraints obtained from stable

isotope thermometry of quartz–tourmaline pairs. This

Table 2

Oxygen isotope data

Sample Mineral d 18OSMOW

WW-EM-1 quartz 12.37

WW-EM-2 tourmaline 9.52

WW-EM-3 tourmaline 10.30

WW-EM-4 scheelite 6.32

WW-EM-5 scheelite 6.07

D. Marshall et al. / Chemical Geology 194 (2003) 187–199196

intersection defines an area ranging from 365 to 498

jC, with corresponding pressures along the isochoric

paths ranging from 700 to 2250 bars. These pressures

correspond to burial depths ranging from just over 2

to approximately 7 km. This is in reasonably good

agreement with the Cretaceous metamorphic grade in

the area of upper greenschist to lower amphibolite.

The P–T data obtained from the emerald veins is also

somewhat ambiguous as to the source of the emerald.

The higher end of the temperature range is probably

slightly higher than the regional metamorphism and

suggests some igneous input of at least heat and

possibly Be, B and Si. However, the lower end of

the P–T data is in good agreement with the regional

metamorphism with a possible distal fluid (Be, B and

Si) source with transport along the Tintina Fault

System.

The state of the fluid during emerald precipitation

is consistent with a single-phase (non-boiling) fluid.

This is shown by the consistent phase ratios within the

fluid inclusions and that the P–T constraints fall

within the one-phase field (Takenouchi and Kennedy,

1964) on a CO2–H2O diagram approximating the

fluid inclusion compositions obtained in this study

(Fig. 8).

8. Discussion and conclusions

The emerald at the Crown showing is the only

occurrence of Cr-bearing emerald in the Canadian

Cordillera and its discovery indicates that there may

be a new exploration target in this area. Future

exploration will be predicated on what is known about

the Crown showing and how it ‘‘pigeon holes’’ into

the existing models of emerald formation. The emer-

ald is found in close proximity to S-type Cretaceous

granite and hosted within mafic schists of sufficient

metamorphic grade to account for a hydrothermal

origin for the emerald. H and O isotope studies (Groat

et al., in press) on the Crown emerald are ambiguous

in delineating a magmatic versus a tectonic–hydro-

thermal origin for the emerald. Oxygen isotope studies

presented here indicate a polythermal paragenesis,

with local isotopic equilibrium attained at the hand

specimen scale. This is consistent with the petro-

graphic and field relationships, which indicate the

primary vein minerals are quartz and tourmaline with

emerald generally being restricted to the vein alter-

ation haloes comprised of a tourmaline–muscovite

schist. This alteration assemblage is overprinted by a

lower temperature assemblage of chlorite, scheelite,

Fig. 8. Pressure– temperature diagram showing the range of possible P–T conditions for emerald growth at the Crown showing. Temperature

constraints are derived from co-existing quartz and tourmaline oxygen isotope data. Fluid inclusion isochores are derived from the interpolation

of Gehrig’s (1980) experimental data and equations of state (Jacobs and Kerrick, 1981; Bakker, 1999). The two-phase field is interpolated from

the data of Takenouchi and Kennedy (1964) for a binary H2O–CO2 fluid with an X-CO2 of 0.05.

D. Marshall et al. / Chemical Geology 194 (2003) 187–199 197

quartz, lepidocrocite and jarosite, during the waning

stages of the hydrothermal system.

Fluid inclusion studies on the quartz–tourmaline–

emerald veins and the granite suggest markedly differ-

ent fluid compositions within the emerald and granite,

as the granite contains low volatile melt inclusions

and aqueous volatile poor secondary inclusions, while

the primary fluid inclusions in emerald are CO2-

bearing. Although carbonate units have been reported

within the Fire Lake metavolcanic rocks (Fig. 1) on

the eastern side of the granite, no such carbonates

have been observed in the general proximity of the

emerald mineralization, which lies to the west of the

granite. Geochemistry of the nearby Cretaceous gran-

ite (Groat et al., in press) indicates anomalously low

Be values, suggesting beryllium may be sourced from

elsewhere and transported to the emerald showing via

large crustal structures (Nwe and Grundmann, 1990;

Nwe and Morteani, 1993; Grundmann and Morteani,

1989) such as the Tintina Fault system, a major fault

14 km to the southwest of the emerald veins. It is

conceivable that the majority of the Be was taken

from the melt prior to crystallization of the granite.

This would require an elevated partitioning coefficient

for Be between the melt and fluid. Despite the require-

ment of an elevated partitioning coefficient, the Creta-

ceous granite is the most likely source of Be because

the emerald veins are 14 km from Tintina Fault, and

there is no evidence of fluid transport from the

Tintina.

The temperature range of emerald precipitation is

365–498 jC. The upper part of this temperature range

is probably slightly in excess of the regional meta-

morphic grade. This may be an artifact of partial

disequilibrium in the 18O of the quartz and tourmaline,

or it could indicate that the granite contributed to the

overall heat budget during emerald formation. Forma-

tional pressures correspond to burial depths on the

order of 2–7 km. As the emerald is either related to

the intrusion of the granite or peak metamorphism,

both of which occurred at approximately 112 Ma, this

constrains the uplift rate in the area to a very slow rate

of approximately 0.02–0.07 mm/year.

Studies underway include hydrogen and oxygen

isotope studies on the Cretaceous granite, bulk-leach-

ate fluid inclusion analyses and Ar–Ar studies on the

muscovite from the vein alteration haloes. The Ar–Ar

data will only be able to delineate a post-granite

tectonic–hydrothermal for the veins as peak meta-

morphism coincides with granite emplacement. O–H

isotopic studies, on the other hand, may provide a link

between the emerald O–H data and the magmatic

fluid composition and remove the need to rely upon

literature O–H data for S-type granites. Despite the

growing body of scientific data on the Crown emerald

veins, and although the combined fluid inclusion and

stable isotope data favour a tectonic–hydrothermal

origin with probable heat input from the Creatceous

granite, at present, there is still no clear indication of a

tectonic–hydrothermal versus a magmatic or hybrid

origin for these deposits.

Acknowledgements

Financial support for the project was provided by

NSERC grants to DM and LG. Danae Voormeij is

thanked for helping with fluid inclusion measure-

ments, field work and draughting. Special thanks to

True North Gems and Expatriate Resources for field

support during this project. Holly Keyes is thanked for

help with the figures. Anita Lam assisted with some of

the field work. [RR]

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