U-Pb and Pb-Pb age constraints on Paleoproterozoic
magmatism, deformation and gold mineralization in the
Omai area, Guyana Shield
Christine Norcross
a, Donald W. Davis
b,*, Edward T.C. Spooner
a,
Alison Rust
aaDepartment of Geology,Uni6ersity of Toronto,22Russell Street,Toronto,Ont.,Canada M5S3B1 bEarth Science Department,Royal Ontario Museum,100Queen’s Park,Toronto,Ont.,Canada M5S2C6
Received 16 February 1999; accepted 3 December 1999
Abstract
The Omai intrusion-centred Au-quartz vein system, the largest Au producer presently operating in the Guyana Shield, was sampled for detailed U-Pb and Pb-Pb geochronology and petrological investigation. The age of a metavolcanic/sub-volcanic unit in the host rock sequence is 212092 Ma. Zircon analyses from the main body dioritic rocks give U-Pb ages of 209496, 2092911 and 2096+11/ −10 Ma. Magmatic titanite and apatite that grew in hornblende-rich peripheral phases of the intrusion define a consistent Pb-Pb age of 209491 Ma, in agreement with the zircon data. Colourless titanite and rutile from strongly altered phases of the intrusion, along with low-U apatite and feldspar, define a significantly younger Pb-Pb isochron age of 200295 Ma. The igneous ages agree with data from similar units in the Guyana Shield and West Africa, showing that2100 Ma was a time of significant intrusive activity. The ages obtained for the deformed metavolcanic and undeformed intrusion at Omai define a 2692 Ma bracket for Trans-Amazonian deformation in central Guyana. Previous fluid inclusion studies indicate that the mineralizing solutions at Omai were too CO2-rich to form titanite, and the titanite-bearing sample is unmineralized,
suggesting that it was not altered by gold-bearing solutions. Therefore, the 200295 Ma age is interpreted as a late hydrothermal overprint that formed titanite and reset rutile. Zircon and baddeleyite from a thick gabbro dyke of the Avanavero Suite, which cuts the Omai pluton, define an age of 179494 Ma, ruling out the dyke as a source for the late thermal effects. The hydrothermal age may record the passage of fluids released by deep crustal metamorphism due to late-stage tectonic underplating as previously proposed for the Superior province. © 2000 Elsevier Science B.V. All rights reserved.
Keywords:U-Pb; Zircon; Gold; Geochronology; Guyana; Trans-Amazonian
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1. Introduction
The Precambrian of South America contains two large cratonic regions, the Amazonian and
* Corresponding author. Tel.: +1-416-5865811; fax: + 1-416-5865811.
E-mail address:dond@rom.on.ca (D.W. Davis)
C.Norcross et al./Precambrian Research102 (2000) 69 – 86 70
Sa˜o Francisco cratons. The Amazonian craton is further subdivided by the Amazon river basin into the Guyana and Guapore´ Shields (Fig. 1), to-gether covering eastern Venezuela, Guyana, Suri-name, French Guyana, the northeastern part of Brazil and a small eastern section of Colombia. The geochronology of the South American Pre-cambrian is broadly known, and in many areas in the cratons of Brazil it is well constrained by
detailed studies (e.g. Machado et al., 1996a,b; Teixeira et al., 1996). The focus of this work is the Trans-Amazonian orogeny in Guyana. This was a major crust-forming episode that has been dated in the approximate range 1900 – 2200 Ma (Gibbs and Barron, 1993). It is broadly correlative with the Eburnean orogeny in West Africa (Milesi et al. 1992). The nature and timing of Trans-Amazo-nian activity has not been described in detail
Fig. 2. Geology of the Omai intrusion-centred Au-quartz vein hosted deposit. (a) is a plan section at 122 m below surface, showing approximate locations of samples (boxed numbers=OM95-x). Sample OM95-27 is a drill core obtained from beneath the Wenot Lake zone and the location is not shown. (b) is a cross-section of the deposit. (Modified from Bertoni et al., 1991)
across the continent. Were events coincident in time in different areas, or did they occur di-achronously? The large scale of the Trans-Amazo-nian, and its apparent duration of several hundred million years, makes detailed reconstruc-tion of its activity essential to understanding the geologic evolution of South America. The first step in this work is to define periods of specific Trans-Amazonian activity in specific areas. Rela-tively little work has been done in the northern cratonic regions, such as the Guyana Shield, where minimal development and access combined with heavy tropical cover and weathering have made reliable geological sampling difficult. The Omai mine provides an important opportunity to extend coverage of Trans-Amazonian rocks in this area.
The Guyana Shield has been a producer of alluvial and eluvial gold for over a century and, as such, makes an excellent exploration target. Sev-eral deposits are presently being mined in the granite – greenstone terranes. The Omai diorite in-trusion-centred Au-quartz vein system in Guyana
(Fig. 2) is the largest current open-pit gold pro-ducer in the Guyana Shield, with estimated re-serves of four million ounces of gold. This open pit operation, owned by Cambior, Golden Star Resources and the Government of Guyana, pro-vides what is probably the largest hard-rock out-crop in the Guyana Shield at the present time.
C.Norcross et al./Precambrian Research102 (2000) 69 – 86 72
2. Background
The Guyana Shield can be divided into four principal Precambrian terranes: (inset, Fig. 1) the
Archean Imataca Complex, Paleoproterozoic
greenstone belts, the Uatuma˜ Group and sedi-mentary sequences such as the Roraima forma-tion. The Imataca Complex in northeastern Venezuela includes granulite gneiss terranes, iron
formations and metasediments. This
al-lochthonous unit is thought to be at least 3.4 Ga old, and suffered major deformational events at
:2.7 and 2.0 Ga (Wirth et al., 1990).
The first major continental crustal development in the Shield occurred during the early
Protero-zoic at :2.3 – 2.1 Ga. This created a series of
greenstone belts and associated gneisses and am-phibolites that are similar to Archean granite – greenstone complexes found in shield regions around the world. The greenstone sequence in the Guyana Shield generally changes from low-K basalts through intermediate and felsic volcanics to volcanic and chemical sediments. Most of the volcanism is thought to be of submarine origin from multiple centres (Gibbs and Barron, 1993). Greenstone belts across the Guyana Shield in-clude the Pastora group in Venezuela, the Barama-Mazaruni group in Guyana, the Marowi-jne group in Suriname, and the Maroni group in French Guyana.
Following volcanism and associated plutonism, all the existing crustal fragments were assembled during the Trans-Amazonian orogeny. This tec-tonothermal episode was originally defined by Hurley et al. (1967), based on a large cluster of K-Ar and Rb-Sr radiometric ages around 2000 Ma. Since then, it has been found that throughout the cratons of the continent there is an abundance of U-Pb, Rb-Sr, K-Ar and Sm-Nd radiometric data that cluster in the range 1900 – 2200 Ma, indicating that this period was a time of signifi-cant deformation, metamorphic and intrusive ac-tivity, followed by crustal cooling (e.g. Cordani and de Brito Neves, 1982; Gibbs and Barron, 1993). A limited number of reliable geochronolog-ical studies of the granite – greenstone terranes of the Guyana Shield have been published to date and are summarized in Table 1. Ages range from
1850 to 2350 Ma for metavolcanics and 1900 to 2250 Ma for syn- to post-orogenic plutons, in-cluding errors. As in other Trans-Amazonian ter-rains, the Rb-Sr and K-Ar ages form a cluster around 2000 Ma. However, Rb-Sr and K-Ar methods are generally considered unreliable to date crystallization ages due to the likelihood of isotopic resetting during late stages of the Trans-Amazonian orogeny.
Within the Trans-Amazonian period of activity, two major stages of intrusion can usually be recognized. The first stage produced pre- and syntectonic intrusions, occasionally associated with greenstone belt volcanism. In the Guyana Shield, these rocks were affected by cataclastic deformation in WNW and ENE directions (Gibbs and Barron, 1993). Following the final stages of deformation, a second phase of intrusive activity created more potassic granitic rocks and other intrusions ranging from quartz syenite and diorite to tonalite in composition. In Venezuela and northern Guyana these are termed ‘younger gran-ites’ and are thought to have followed the main Trans-Amazonian deformation because they do not show the same deformational characteristics as their host metavolcanics (Gibbs and Barron, 1993).
The mid-Proterozoic saw the formation of the Uatuma˜ group of felsic volcanics and granitoid intrusions from 1.7 to 1.9 Ga, followed by the development of sedimentary sequences, such as the Roraima formation. These sequences were later intruded by mafic dykes such as the Avanavero suite, formerly known as the Roraima Intrusive Suite (e.g. Gibbs and Barron, 1993; Sid-der and Mendoza, 1995).
3. Geology and mineralization
The Omai Au-quartz vein system is located at
58° 45%W and 5° 28%N on the west bank of the
thought to be mostly of submarine origin based on textural, chemical and mineralogical evidence (Gibbs, 1987). In the vicinity of the Omai deposit, the metavolcanics have undergone regional ductile deformation and greenschist facies metamor-phism. The Omai pluton is a lobate intrusion of dioritic affinity, with margins that clearly cross-cut the ductile deformational fabric in the host metavolcanics. The margins of the Omai pluton are marked by hornblende-rich diorite and horn-blendite rocks, whereas the main body ranges in composition from diorite to quartz diorite (Fig. 2). Hydrothermal alteration is pervasive through-out the intrusion, but varies in intensity with location. Gold is mainly found in the quartz diorite. The marginal hornblende-rich phases and metavolcanics are rarely mineralized. However, a zone of gold enrichment in saprolite above pri-mary gold mineralization, termed the Wenot Lake
zone, is located in metavolcanics about 500 m south of the intrusion, and is presently being mined.
Gold, as free grains of the native metal, is hosted largely by quartz veins disseminated throughout the main body of the intrusion, but can also be found as micro-inclusions in pyrite within the wall rock (Bertoni et al., 1991). The auriferous vein stockwork consists mainly of quartz, ferroan carbonate, sulphides and scheelite, ranging in size from stringers up to 5 cm in width
to occasional larger (1 m) veins. The Omai
deposit also contains auriferous tellurides and
bis-muthinides. Hydrothermal alteration shows
phases that are spatially and temporally related to gold mineralization. These phases include carbon-ate alteration, sulphidation and silicification. Car-bonate alteration resulted in the formation of ferroan carbonate both in veins and wall rock
Table 1
Ages of greenstone belts and intrusions in the Guiana Shield
Rock type Location Age, Ma Method Reference
Pastora group Day et al., 1995
Metavolcanic 2131910 U-Pb, zircon
(Venezuela)
Metagraywacke Barama-Mazaruni 22509106, U-Pb, zircon Gibbs and Olszewski, 1982 2244943
group (Guyana)
Priem et al., 1982 Metavolcanics Marowijne group 19509150 Rb-Sr
(Suriname)
Gruau et al., 1985 Metavolcanics Paramaca series 2210990 Sm-Nd
(French Guiana)
U-Pb, zircon
2227939 Gibbs and Olszewski, 1982 Barama-Mazaruni
Bartica gneiss
group (Guyana)
Lerouge et al., 1996. Pb-Pb, zircon
2123911, Pegmatite, yaou Paramaca series
(French Guiana)
granite 2127910
Voicu et al., 1997 Sm-Nd
P. Klipfel, personal communication, Pastora group
KM24 granite 2087921 U-Pb, zircon
November 4, 1998. (Venezuela)
Paramaca series,
Granites 2030965, Rb-Sr Pb-Pb K- Teixeira et al., 1996 from Gibbs and Bar-ron, 1993.
2083939, Ar (model) (French Guiana)
2032961 Barama-Mazaruni 2015980
Younger granitea K-Ar Snelling and McConnell, 1969
group (Guyana)
Younger granitea Barama-Mazaruni 19459100 K-Ar Williams et al., 1967
group (?) (Guyana) Marowijne group (?)
Granitoids and acid 1810940 Rb-Sr Priem et al., 1971 volcanics (Suriname)
C.Norcross et al./Precambrian Research102 (2000) 69 – 86 74
Fig. 3. U-Pb concordia diagram of zircon analyses from sample OM95-27, the felsic metavolcanic. Numbers refer to analyses in Table 2. Error ellipses are 2s.
on several characteristics, as follows. Attempts were made to select zircons that were clear, crack-free and lacking inclusions, cores and over-growths. However, the poor quality of samples often made selection of cracked and cloudy grains necessary. Titanite fragments were divided based on colour and selected for internal clarity and lack of cracking. Apatite grains are prismatic whereas the feldspars are generally anhedral. Both were selected for clarity and lack of cracking. Rutiles were chosen for crystal shape, colour and luster (an indicator of freshness). Rutile aggre-gates were washed in HF for 15 min in an ultra-sonic cleaner to remove surrounding silicate minerals. Selected zircons were abraded (Krogh, 1982), but many were left unabraded due to the possibility of shattering because of internal cracks. No other minerals were abraded. Large titanite, apatite and feldspar fractions were weighed, whereas weights for zircon, baddeleyite and rutile fractions were estimated by eye and are
likely accurate to about950%. Zircon,
baddeley-ite and rutile were digested in bombs, while titan-ite, apatite and feldspar were digested in Savillex capsules. For zircon, U and Pb were separated
using HCl with 50 ml anion exchange columns
following the method of Krogh (1973). Titanite, apatite, rutile and feldspar were passed through
500 ml anion exchange columns with HBr,
follow-ing the method of Corfu (1988). Pb blanks are 1 pg for small column and 2 pg for large column chemistry. U blanks are taken to be 0.1 pg. U and
Pb were loaded onto Re filaments using H3PO4
and silica gel. Isotopic analysis was carried out on a VG354 mass spectrometer in peak jumping mode, with either a Faraday collector (for large signal samples) or Daly detector (for small signal samples). The mass discrimination correction for the Daly detector is 0.40% per AMU and the thermal mass discrimination factor is 0.10% per AMU. Common lead in the apatite, titanite and rutile samples made ages variably dependent on the isotopic composition defined for the initial common Pb. Data from the feldspar from OM95-2 were used as a common lead correction on all regressions for these minerals. Data are plotted with two sigma error ellipses on U-Pb and Pb-Pb diagrams (Figs. 3 – 9). Regressions are calculated during pre-gold and gold-forming stages.
Sul-phides are mostly pyrite, but also include minor sphalerite, galena, and chalcopyrite. Pyrite miner-alization occurred before, during and after gold deposition, as did silicification. Increased sulphi-dation and silicification correlate with increased veining and gold mineralization.
A several hundred metres thick gabbro dyke, believed to be a member of the Avanavero Suite,
cuts the mineralized pluton at a depth of :200 –
300 m (Bertoni et al., 1991). Members of this group of mafic rocks occur throughout the Guyana Shield in the form of sills, dykes and other irregular bodies and have been dated by K-Ar, Rb-Sr and Ar-Ar methods. Sidder and Mendoza (1995) estimated the age at between 1650 and 1850 Ma, based on compiled data from several sources.
4. Analytical methods
Sample sizes ranged from 6 to 25 kg before crushing. Heavy mineral separates from the
Wi-lfley table were screened to −70 mesh, and then
Fig. 4. Pb-Pb diagram showing feldspar analyses from OM95-2, and hydrothermal titanite, apatite and rutile analyses from OM95-2, OM95-26 and OM95-28, respectively. The error el-lipses are highlighted by circles. The line represents a regres-sion of all analyses except 7 (titanite) and 29 (feldspar). The Stacey and Kramers (1975) growth curve is also shown.
using the program of Davis (1982). Resulting age errors are quoted at 95% confidence levels.
5. Samples and results
Mineralized and non-mineralized samples,
showing varying degrees of hydrothermal alter-ation, were chosen in order to obtain ages on both igneous emplacement and hydrothermal crystallization (Fig. 2). U-Pb analytical data for samples with highly to moderately radiogenic Pb are presented in Table 2. Pb-Pb data for low
U/Pb titanite, apatite, rutile and feldspar fractions
are presented in Table 3. Note that analyses 5 – 6 are given in both tables.
5.1. OM95-27 Meta6olcanic host rock
Sample OM95-27 is an intermediate-felsic
metavolcanic/subvolcanic rock recovered as drill
core from 130 m depth under the Wenot Lake
zone. It consists of plagioclase and minor quartz phenocrysts in a fine-grained quartz-feldspar ma-trix. Secondary carbonate, chlorite and sericite are abundant, with minor pyrite and accessory zircon.
Zircons from this sample are elongate (l:w
3:1), clear and colourless, with an average length
of :50 mm. Two abraded grains were run as
separate analyses, along with one unabraded piece that contained longitudinal cracks. Data from the abraded fractions are concordant and, with the slightly discordant datum from the unabraded
fraction, define an age of 212092 Ma (Fig. 3).
5.2. OM95-2 Marginal hornblende diorite
Three samples of hornblende-rich marginal phases were collected. OM95-2 is a porphyritic hornblende diorite taken from the north edge of the pit. The hornblendes in this unit are stubby crystals showing a sub-parallel orientation. The matrix consists of quartz and plagioclase with minor smaller amphiboles and homogeneously distributed magnetite crystals. Strong hydrother-mal alteration produced carbonate, sericite and epidote. Alteration of the hornblendes to chlorite and actinolite is pervasive. Accessory phases
C
U-Pb data for zircon, baddeleyite, titanite and apatitea
Th/U Pbcom(pg) 207Pb/ 206Pb/238U 207Pb/235U 207Pb/206Pb Age
Weight (mg) %Disc
Number Fraction U (ppm)
(Ma)
204Pb
OM95-27meta6olcanic
746.9 0.389498 7.068920
0.002 98 2120.092.2 0
1 Ab z, flat
1 0.28 0.9
355.4 0.3891912
2 1 Ab z, round 0.001 84 0.29 0.8 7.060928 2119.694.2 0.1
669.7 0.379498 6.877918 2117.892.2 2.5 5.0
1 z, frag, cracked 0.006 169 0.6 3
OM95-2hornblende diorite
165.5 0.381698 6.820918
4 Brown titanite 0.477 61 3.83 560 2092.991.8 0.5
49.38 0.3580913 6.096929 2007.595.1 2.0 36
5 colourless titanite 0.050 9.5 3.13
1.69 266 32.70 0.352098 5.965928 1998.696.3 3.2 6 colourless titanite 0.290 6.3
43.09 0.365298 6.365924 2048.394.4 2.4 243
11
7 colourless titanite 0.226 2.28
OM95-4bladed hornblendite
79.01 0.384499 6.875921
8 Brown titanite 0.303 79 2.53 1101 2094.292.7 -0.1
417.7 0.381298 6.822918 2095.491.6 0.8 517
9 Dark brown titanite 0.181 397 1.48
179
Colourless titanite 0.288 11 2.12 65.77 0.376499 6.730923 2093.993.0 2.0 10
OM95-11quartz diorite
1.64 12.1 67.75 0.3279910 5.843924 2087.694.2 14.3 0.001
12 3 Ab z, fuzzy core 228
180.9 0.313498 5.562926
13 2 z, cracked, fuzzy core 0.003 230 1.3 10.8 2080.797.0 17.7
174.6 0.290096 5.148924 2081.197.4 23.9
16 3 z, cracked 0.005 308
OM95-28quartz diorite
31.02 0.3095914 5.512944 2087911 19.0
20 4 z, cracked 0.003 304
OM95-1normal diorite
21 4 Ab z, tiny 0.001 267 1.78 10.2 93.58 0.3547910 6.297922 2081.193.6 6.9
OM95-3gabbro dyke
25 1 Baddeleyite 0.001 103 0.05 0.5 485.1 0.3166914 4.766923 1786.094.5 0.8 2099 0.3187910 4.808915 1790.092.5 0.4
aNotes: Errors are two standard deviations. Pb blank–1 pg (zircon), 2 pg (titanite and apatite); U blank–0.1 pg. Ab, abraded, z, zircon. Pb
com, Common Pb,
including blank; calculated using blank isotopic composition. Th/U calculated from radiogenic 208Pb/206Pb ratio and 207Pb/206Pb age assuming concordance.
Fig. 6. U-Pb concordia diagram showing OM95-11 zircon analyses. The two sigma error ellipses are highlighted by circles. The line represents a regression of all data except analysis 16. Only analysis 12 was from an abraded fraction.
Fig. 8. U-Pb concordia diagram showing zircon analyses from OM95-1 (Diorite). The two sigma error ellipses are highlighted by circles. The upper line (2096911 Ma) was regressed with-out analysis 21 (the only analysis from an abraded fraction). The lower line (211096 Ma) is a regression through all four data.
Fig. 7. U-Pb concordia diagram showing zircon analyses from OM95-28 (vein wall rock). The two sigma error ellipses are highlighted by circles. The line represents a regression of all data. Analyses were from abraded fractions except for 20.
Fig. 9. U-Pb concordia plot showing analyses of zircon and baddeleyite from a late mafic dike. The line is a regression through the three data.
C.Norcross et al./Precambrian Research102 (2000) 69 – 86 78
Table 3
Pb isotopic analyses for secondary minerals from hydrothermally altered rocks at Omaia
Fraction
Number Weight (mg) U (ppm) PbCom (pg) Corr. Coef. 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb
OM95-2Hornblende Diorite Colourless
5 0.050 9.5 36 0.9989 306.0916.4 51.292.0 295.4914.8
titanite
Colourless 0.290 6.3 266
6 0.9491 157.7491.05 32.8190.15 104.0190.76
titanite
7 Colourless 0.226 11.0 243 0.9744 236.8291.78 43.3090.25 179.591.3 titanite
0.553 0.16 3180 0.9781
28 Feldspar 15.682 15.356 35.5590.12
90.036 90.024
0.29 31 7415 0.8606
29 Feldspar 17.595 15.491 40.3090.16
90.040 90.034
OM95-26Quartz Diorite 0.175
30 Apatite 1.6 3224 0.9838 17.010 15.518 36.4490.11
90.036 90.027
OM95-28Quartz Diorite
0.048 4.8
31 Rutile, dark 766 0.9338 21.430 16.042 54.3090.19
90.038 90.039
0.402 6.6 5042 0.9693
Rutile, dark 25.981 16.635
32 82.7090.26
90.039 90.043
16.694
0.683 24.0 30426
Rutile, yel- 0.9971 26.465
33 64.4590.23
90.041
low 90.038
aFootnotes to Table 2 apply.
Two fractions of plagioclase feldspar were analysed from this sample, to be used for com-mon lead corrections. Fig. 4 (insert) shows the Pb-Pb plot of the two points, with the Stacey and Kramers (1975) curve for reference. Analysis 28 (Table 3) contains the most primitive Pb and plots just above the curve, whereas analysis 29 plots
below it but shows much more evolved 206
Pb/
204Pb. The more radiogenic Pb data from other
minerals plot colinearly with datum 28, but do not regress successfully with the more evolved feldspar datum 29. The feldspar fractions were picked to be as fresh as possible but slight turbid-ity was evident in many of the grains, indicating the presence of some alteration. The high U con-tent of analysis 29 (Table 3) suggests that it was from an impure feldspar fraction that may not
have preserved a primary Pb isotopic
composition.
Analysis of one fraction of brown titanite gave
a data point only 0.5% discordant with a 207Pb
/
206Pb age of 209392 Ma when corrected for the
initial composition of the feldspar (analysis 4, Fig. 5). The brown titanite is relatively high in U and its Pb is fairly radiogenic. Three fractions of colourless titanite fragments were also analysed from this sample. These contained much lower U concentrations and had less radiogenic Pb. Using the feldspar datum to correct for common Pb, the data points are 2 – 3% discordant and show
vari-able 207
Pb/206
Pb ages of 204894, 200895 and
5.3. OM95-26 Altered hornblendite
OM95-26 is a hornblendite with strong hy-drothermal alteration chosen from the southeast hornblende-rich margin of the pluton. It contains bladed amphibole crystals that are pseudo-morphed by fine-grained carbonate, sericite, chlo-rite and some minor plagioclase and opaque minerals including pyrite (with microscopic gold inclusions), magnetite and chalcopyrite. The ma-trix consists mostly of quartz and plagioclase, usually in a granophyric intergrowth. The quartz
is filled with H2OCO2 fluid inclusions and the
plagioclase is strongly altered to fine-grained sericite and calcite. Apatite crystals were found in the matrix, as were fine rutile needles and pleochroic haloes indicating possible zircons.
Apatite was the only potentially datable min-eral that could be recovered from this sample. A
fraction of well-formed, clear apatites :250 mm
long was analysed. The resulting datum (analysis 30, Fig. 4 insert) is non-radiogenic and plots close to the feldspar data from OM95-2.
5.4. OM95-4 Marginal hornblendite
OM95-4 is a hornblendite taken from the west side of the pit, where field relations show that it crystallized in situ. Large, zoned and twinned hornblende blades that are partly altered to chlor-ite and actinolchlor-ite make up 75% of the sample. The matrix is quartz and albite, occasionally with granophyric texture, and the sample shows mod-erate hydrothermal alteration to carbonate, epi-dote and sericite. Opaque minerals include pyrite, magnetite, chalcopyrite and bornite. Zircon was not seen in thin section. Accessory minerals in-clude large, zoned euhedral titanites, associated with magmatic hornblende that grew in situ in-ward from the intrusive contact, and large euhe-dral apatite. Both appear to be of the same generation as the hornblende.
Fragments of titanite up to 150mm in size were
recovered. Dark red-brown fragments were from the cores of crystals, colourless fragments were from the edges, and mottled, medium-brown frag-ments were presumably from areas between the centres and the edges. One multi-grain fraction of
each colour was analysed (Table 2). Uranium contents increase with depth of colour. One data point is within error of concordia while the others
are only slightly discordant, but the 207Pb/206Pb
ages are essentially the same and average to 209491 Ma (Fig. 5: 57% probability of fit).
One fraction of apatite crystals was also analysed from this sample. Grains chosen for analysis were generally colourless, well-formed
prisms :300 mm long, with slightly pitted
sur-faces. Correcting for initial common Pb using the
feldspar datum from OM95-2 gives a 207
Pb/206
Pb
age of 2089913 Ma. The data point on a
concor-dia concor-diagram is several percent discordant (Fig. 5).
5.5. OM95-11 Southeast quartz diorite
Sample OM95-11 is a quartz diorite with strong hydrothermal alteration from the hanging wall of a major quartz vein system in the south end of the pit. The matrix consists of quartz and plagioclase with chlorite, sericite, carbonate and rutile as alteration minerals. Opaque minerals include pyrite and minor magnetite, with the former often
containing blebs of gold 30mm in size. Zircons
were found in thin section only as very small
grains (30 mm long) in chlorite, and often only
by their pleochroic haloes so it is unclear if the grains seen in thin section are the same generation as those recovered by mineral separation.
Zircons recovered from this sample are smaller than those from the metavolcanic unit, colourless, and often cracked with cloudy interiors. Five fractions were dated, but only one was abraded due to the poor quality of the grains. The abra-sion did not improve concordance. One highly discordant datum that is not colinear with the rest was not used in the regression (Fig. 6). The lead
loss line gives an age of 209496 Ma, with a 50%
probability of fit and a lower intercept close to zero.
5.6. OM95-28 Central quartz diorite
C.Norcross et al./Precambrian Research102 (2000) 69 – 86 80
The vein material directly adjacent to the sample
contains visible gold. Zircons about 50mm in size
were found in thin section. OM95-28 also con-tained rutile crystals large enough to be recovered. In thin section rutile is usually found as clusters of striated, twinned crystals, dark brown to reddish brown in colour, inside or bordering the matrix quartz and plagioclase.
Four fractions of colourless, slightly cracked zircons, three of which were abraded, were analysed. All data are highly discordant, with the unabraded fraction giving the most discordant datum, which is not quite collinear with the other three. The three most concordant data were used to plot a Pb-loss line. The upper intercept with
concordia is 2092911 Ma, with a 36%
probabil-ity of fit and a near-zero lower intercept age (Fig. 7).
Rutile was recovered from this sample as small (0.2 – 0.25 mm) bundles consisting of brown elon-gate crystals, some of which show elbow twinning. Rutile has previously been used to provide mini-mum age constraints on mineralizing or signifi-cant hydrothermal events (e.g. Davis et al., 1994b). Three fractions were separated for analy-sis. The freshest looking consisted of ten clusters of euhedral dark brown crystals with shiny stri-ated faces. A second fraction consisted of a large number of similar clusters of brown crystals that were less shiny and euhedral than the first frac-tion. The third fraction consisted of anhedral yellow-brown rutile. Pb from the rutile is only moderately radiogenic. On a U-Pb concordia dia-gram, the data points are very imprecise because of the small proportion of radiogenic Pb. On a Pb-Pb diagram the three analyses (31 – 33, Table
3) define an age of 2020956 Ma when regressed
together with the most primitive feldspar datum (Fig. 4, insert).
5.7. OM95-1 North diorite
OM95-1, is an altered diorite from the north end of the pit. It consists mostly of sericitized plagioclase, quartz, chlorite and epidote. Opaque minerals include magnetite and pyrite, with no observable gold. The zircons found in thin section are similar to those in OM95-11 (i.e. they are
extremely small and were seen in thin section only when they formed a pleochroic halo in chlorite).
Zircons recovered from this sample are mostly small, cloudy and cracked. Discordant data from the three unabraded fractions define a Pb-loss line
with an upper intercept of 2096 + / −11 Ma (Fig.
8: 13% probability of fit). The abraded fraction gives the most concordant datum but is not colin-ear with the three other data points. When re-gressed with the two most discordant points, this
datum defined an age of 211096 Ma (Figs. 8 and
73% probability of fit). The lower intercepts for the regressions are between 500 and 600 Ma.
Abrasion of cracked and altered zircons can produce scattered or even reversely discordant data, possibly due to within-grain Pb and U mo-bility in altered domains followed by U-Pb frac-tionation when part of the grain is removed by abrasion (Davis et al., 1982). Therefore, the inter-cept age determined from the three unabraded fractions is considered to be the most reliable in spite of its lower probability of fit.
5.8. OM95-3 Gabbro dyke
Sample OM95-3 is a drill core sample of the gabbroic dyke that cuts through the Omai pluton at depth. The sample was obtained from beneath the main pit. In thin section no dateable minerals were seen, although a few pleochroic haloes were observed, indicating the possible presence of zircon.
Mineral separation yielded only a few zircon and baddeleyite grains. The zircon mostly consists of cracked and altered elongate grains, whereas the baddeleyite grains are very tiny. Two zircons were analysed from this sample, a crack-free grain that was abraded prior to analysis, and a second elongate, prismatic, cracked crystal that was not abraded. One flat, fresh-looking grain of badde-leyite was also analysed. The data from the un-cracked abraded zircon and the baddeleyite (Table 2) are nearly concordant, whereas the da-tum from a cracked unabraded zircon (analysis 27) is reversely discordant (Fig. 9). The regression
of the three data points defines an age of 179494
Reversely discordant data have only rarely been observed from zircon. Lack of equilibration be-tween spike and sample can produce spurious reversely discordant data, but this is unlikely to be
the case with analysis 27 because its 207Pb/206Pb
age is older than those of the near concordant data and the three data define a statistically con-sistent regression with a fairly old lower concordia intercept age. The reversely discordant grain showed a relatively high U concentration, and lower concordia intercepts near 1000 Ma are typi-cal for high U zircons from mafic rocks, probably because they begin to accumulate radiation dam-age-induced Pb loss earlier than low U zircons. It appears that this may be a real case where U loss from zircon exceeded Pb loss. Unfortunately, the few remaining zircon grains from this sample are quite altered and, since they probably underwent pronounced secondary Pb loss, they are likely to give complex discordia that would not clarify the cause of the reverse discordance.
6. Discussion
6.1. Ages of magmatism and deformation
The age of one metavolcanic rock in the Omai
area is defined by this work to be 212092 Ma
(Fig. 3). The most precise age obtained for metavolcanics in the Guyana Shield before this study was on the Pastora Group in Venezuela, less than 500 km from the Omai location (Day et
al., 1995). This unit gave an age of 2131910 Ma,
in agreement with the metavolcanic age from Omai. This suggests that greenstone belts in the Guyana Shield may have formed over a short time span despite having been derived from differ-ent volcanic cdiffer-entres. However, U-Pb dating by Gibbs and Olszewski (1982) on the Barama-Mazaruni assemblage and Sm-Nd work by Gruau et al. (1985) on the Paramaca series in French Guyana defined ages for metavolcanics of 22509
106 and 2210990 Ma, respectively. These ages
are significantly older than in this work although their large errors make it difficult to define the total age span of volcanism.
The five samples obtained from the Omai plu-ton are representative of all the main intrusive phases. The discordance of much of the zircon data results in relatively large errors (6 – 10 Ma). However, the agreement of zircon results from
three samples (209496, 2092911 and 2096+11/
−10 Ma) indicates that the ages obtained are
probably reliable indicators of primary crystalliza-tion. The igneous titanite and apatite data give an
average 207Pb/206Pb age of 209491 Ma (Fig. 5:
29% probability of fit, corrected using the most primitive feldspar datum). This age agrees well with the less precise zircon ages and is probably the best estimate for crystallization of the pluton, particularly because textural evidence indicates that the titanite grew as part of the magmatic assemblage. The only known similar U-Pb age for
a granitoid in the Guyana Shield is a 2087921
Ma zircon age from a granite in a granite – green-stone terrane south of El Dorado, Venezuela (P. Klipfel, Placer Dome Exploration, personal com-munication, November 4, 1998; Table 1).
The ages obtained in this study for the
metavol-canic sample (212092 Ma) and the intrusion
(209491 Ma) fall within the accepted period of
the Trans-Amazonian tectonothermal episode (1900 – 2200 Ma). The metavolcanic units were affected by one or more of the Trans-Amazonian deformational episodes. The Omai pluton does not show the deformation seen in the surrounding metavolcanic units and clearly cross-cuts their steep tilting and ductile strain fabric. The contact of the pluton margin with the metavolcanics is equally undeformed, so the pluton must have been intruded as a member of the Younger Gran-ite group of intrusions after the end of the last Trans-Amazonian ductile deformational episode. Thus, the ages obtained for the metavolcanic unit and the intrusion bracket the time period for Trans-Amazonian deformation in this region of
Central Guyana to 2692 Ma.
6.2. Pb loss from zircon
circu-C.Norcross et al./Precambrian Research102 (2000) 69 – 86 82
lation does not necessarily result in Pb-loss due to leaching. The zircons from the metavolcanic ple are fresher than those from the plutonic sam-ples, which are generally cracked, cloudy, and small. The pluton-derived zircons may be more zoned in U than the metavolcanic zircons. Radia-tion damage in high U zones causes an increase in crystal volume, cracking the crystals and allowing penetration of water and subsequent alteration (Krogh and Davis, 1974, 1975). In contrast, the lack of cracks in the metavolcanic zircons may have prevented water from entering and altering the crystals thus reducing Pb loss from the interi-ors of these grains.
The ca. 500 Ma lower concordia intercept age from OM95-1 probably records an average of multiple or prolonged ancient Pb loss events. Its ambiguous upper intercept age interpretation is probably due to the fact that data from different grains lie on discordia with slightly different lower intercepts. In contrast, regressions of zircon analyses from two other Omai intrusion samples, OM95-11 and OM95-28, have near-zero age lower concordia intercepts. The near-zero lower inter-cept age from these samples indicates that, while they may have undergone ancient alteration and Pb-loss as shown in sample OM95-1, Holocene tropical weathering may have caused extensive leaching of Pb from altered zones in the zircons,
resetting them to give consistent primary 207Pb
/
206Pb ages, despite increasing the discordance of
their data. Thus, paradoxically, the influence of recent weathering may have improved recovery of primary age information by erasing any previous history of complex lead loss. If so, laboratory hydrothermal leaching of altered domains could potentially be used to achieve the same effect.
6.3. Hydrothermal mineral ages
Data on suspected hydrothermal mineral frac-tions from strongly altered samples, with the ex-ception of one colourless titanite fraction that may have contained some of the magmatic phase (analysis 7), regress within error of an isochron on a Pb-Pb diagram (Fig. 4) and define an age of
200295 Ma (29% probability of fit). This age
agrees with a less precise Sm-Nd age of 19959
140 Ma on scheelite from gold-quartz veins at the Omai mine (Voicu et al., 1997). The thermal closure temperature for titanite is in excess of 500°C (Mezger et al. 1991), whereas that of rutile is about 400°C (Mezger et al. 1989). Therefore, the relatively young age defined by these minerals is unlikely to be due to diffusional Pb loss during regional metamorphism, which did not go above greenschist facies.
Although the precision of the isochron age may not reflect its accuracy, the relative radiogenicity of the titanites and the apparent agreement of their Pb-Pb systematics with a variety of other minerals indicates that some hydrothermal event affected the rocks substantially later (ca. 100 mil-lion years) than intrusion of the Omai pluton and
the end of ductile deformation at 209491 Ma.
This event is clearly much older than
emplace-ment of the mafic dike at 179494 Ma. The dike
age shows that the Avanavero Suite is probably related to a thermal event that occurred long after the Trans-Amazonian orogeny.
If the 200295 Ma age dates gold, then
miner-alization would be unrelated to local igneous events. Such a case has been argued for anoma-lously young titanite, rutile, muscovite and scheelite ages in Superior province gold deposits (e.g. Jemielita et al., 1990; Wong et al., 1991; Haynes et al., 1992; Anglin et al., 1996). However, the interpretation of these ages is controversial (e.g. Kerrich, 1994). An example of a greenstone-hosted Au-quartz vein system where rutile gave ages younger than the estimated age of mineral-ization is the Kerr Addison-Chesterville system in the Superior province. At Kerr Addison, rutile in dykes gave relatively young ages of 2630 – 2580 Ma, but mutual cross-cutting relationships be-tween the veins that host the mineralization and a system of ‘albitite’ dykes that contain the rutile indicate that mineralization at that time was not possible, as the age of the dykes is unlikely to be younger than 2670 Ma (Spooner and Barrie, 1993).
case, the age defined by the rutile would reflect the time of recrystallization and be a minimum esti-mate for gold. Fluid inclusion data from mineral-ized quartz veins at Omai indicate that the
mineralizing solutions were of a H2OCO2, low to
moderate salinity composition (Rust, 1996). Flu-ids of this type generally do not form titanite at
moderate temperatures since any Ca2+ in the
system would be used by the excess CO2 in
solu-tion to form carbonate minerals, with the Ti4+
being left to form rutile (e.g. Clark et al., 1986). Thus the secondary titanite, whose data largely control the age interpretation, could not have formed from the same fluids as the gold-associ-ated rutile. This suggests that the secondary titan-ite was formed by an overprint that recrystallized the rutile. This hydrothermal event appears not to have affected all of the Omai samples. For exam-ple, OM95-4 apparently contains only one genera-tion of colourless titanite, which gave an igneous age. A fluid flow episode would be most likely to generate the variable effects seen in the Omai rocks. In any case, the data constrain the age of
gold mineralization to between 200295 and
209491 Ma (the intrusion age of the Omai
plu-ton). The question of how the late hydrothermal event might fit into the context of Trans-Amazo-nian crustal development is discussed below.
6.4. Correlations with West Africa and comparisons with the Superior pro6ince
Since the pioneering work of Hurley et al., (1967) based on early radiometric dating, the Guyana Shield and West African cratons have been widely accepted as having originally been part of a single unit (e.g. Marcoux and Mile´si, 1993). Considerably more geochronology and mapping have since been done in West Africa, yet there is still controversy surrounding the timing of the orogeny that generated, deformed and meta-morphosed ca. 2.1 – 2.2 Ga greenstone belts and plutons. In West Africa this orogeny is known as the Eburnean, and is believed to correlate with the Trans-Amazonian of South America. Granitoid intrusions in West Africa are often divided into ‘belt-type’ and ‘basin-type’, depending on whether they intrude greenstone belts or metasedimentary
‘basins’. Most belt-type plutons in Ghana, for example, give dates around 2170 Ma, whereas the basin-type plutons are younger, in the range of 2120 to 2090 Ma (Hirdes et al., 1992; Davis et al., 1994a). Hirdes et al. (1996) separated the West African region into eastern and western sub-provinces. The eastern subprovince is said to
in-clude slightly older volcanics (2185 – 2150 Ma)
than the western subprovince (2105 Ma). The
2120 Ma age of the volcanics at Omai falls be-tween the ages of these two subprovinces.
Similar to descriptions of the Trans-Amazonian orogeny, descriptions of the timing of the Eburnean orogeny vary considerably with loca-tion. The entire event is considered to have oc-curred broadly between 2200 and 1980 Ma (e.g. Sidder and Mendoza, 1995). In the sedimentary basins of Ghana, Eburnean plutonism, deforma-tion and metamorphism have been more tightly constrained to between 2120 and 2080 Ma (Oberthu¨r et al., 1998). In southern Mali, Lie´geois et al. (1991) constrained Eburnean deformation by dating volcanic and granodiorite rocks to
be-tween 209895 Ma and 2074+9/ −8 Ma,
respec-tively. This age spread is similar to that found in Guyana (ca. 25 million years) although the vol-canic and plutonic rocks dated in Mali are some-what younger.
Gold-associated rocks in Ghana, West Africa, show similar ages to those at Omai. Gold-bearing granitoid intrusions in the Ashanti gold camp have been dated at 2105 Ma, and the late-tectonic
Banso granitoid in the Ashanti belt (209792 Ma)
is coeval within error with the Omai intrusion (Oberthu¨r et al., 1998). In Ghana, gold can be bracketed in the range 2098 – 2105 Ma, based on ages of hydrothermal rutile (Oberthu¨r et al., 1998).
In general, the Eburnean/Trans-Amazonian
C.Norcross et al./Precambrian Research102 (2000) 69 – 86 84
Both crust-forming episodes resulted in juxtaposi-tion of greenstone and metasedimentary terrains. These patterns are likely the result of accretionary orogenies that operated at different times but underwent broadly similar processes, as discussed by Davis et al. (1994a). The analogy with late Archean processes is reinforced by ages of rutile and other hydrothermal minerals in gold deposits of the late Archean Superior province which, as at Omai, are found to post-date their host green-stones by up to 100 Ma (e.g. Jemielita et al., 1990; Wong et al., 1991; Anglin et al., 1996), but corre-spond to ages from less robust geochronological systems such as Rb-Sr and K-Ar ages on miner-als. Krogh (1993) presented evidence that deep crustal metamorphism resulting from late tectonic underplating of the Superior craton caused an influx of hot fluids that were focused along shear zones in the upper crust. In some cases, these may have deposited gold, while in others, pre-existing gold deposits may have been overprinted by these events. If so, similar processes may have affected the Amazonian-West African shield much later, during an analogous stage of its development. This suggests that continental crustal growth mechanisms typical of the late Archean were still operating at 2 Ga.
7. Conclusions
The Omai mine area has provided an age of
212092 Ma on a greenstone metavolcanic
sam-ple of the Barama-Mazaruni group, and an age of
209491 Ma for a dioritic pluton of the Younger
Granite group using zircon, titanite and apatite. These ages are in broad agreement with similar
rocks dated elsewhere, using robust
geochronometers, in the Guyana Shield and in West Africa.
Despite Pb-loss from poor quality zircons in the plutonic rock samples, it was possible to ob-tain consistent dates with reasonable errors from all samples that yielded zircon, possibly due to the intense recent tropical weathering that leached and reset altered domains in the zircons.
The age range of 2692 Ma that separates the
volcanic and intrusive activity at the Omai site
brackets the age of Trans-Amazonian
deforma-tion to between 2120 and 2094 Ma in this
region of central Guyana.
The age of Au mineralization at the Omai
deposit can be bracketed between 200295 Ma
(the age of hydrothermal rutile and titanite) and
209491 Ma (the age of crystallization of the
intrusion). The nature of the 200295 Ma age is
still unclear. Geologic, fluid inclusion and petro-graphic evidence suggest that it may have been a later thermal or hydrothermal event unrelated to, but following, the deposition of gold. A gabbro dyke of the Avanavero suite, which cuts the Omai deposit, was dated by zircon and baddeleyite at 179494 Ma, indicating that it is much younger than, and therefore unrelated to, the hydrother-mal events at about 2000 Ma. By analogy with the Archean accretionary orogeny in the Superior province, the late thermal event at Omai may represent overprinting by hydrothermal fluids re-leased during a period of tectonic underplating.
Acknowledgements
Thanks are extended to Omai Gold Mines and Cambior for access to the Omai Mine site. Thanks also to Kim Kwok, Galena Amelina, Raivo Tahiste, Mark Smethurst and Neil Scully for laboratory assistance, and Gabriel Voicu, Marc Bardoux, Robert Cre´peau, Claude Poulin, Bjarne Westin, Rejean Gourde, Desmond Good-man and Kenrick Morris for assistance. Reviews by N. Machado, and J. Claoue-Long and U.G. Cordani improved the manuscript. Research was funded in part by NSERC operation grant OGP0006114 (ETCS).
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