Constraints on the depth of generation and emplacement of a magmatic epidote-bearing quartz diorite pluton in the Coast Plutonic Complex, British Columbia

Abstract

Petrology and P–T estimates indicate that a magmatic epidote-bearing quartz diorite pluton from Mt. Gamsby, Coast Plutonic Complex, British Columbia, was sourced at pressures below ∼1.4 GPa and cooled nearly isobarically at ∼0.9 GPa. The P–T path indicates that the magma was within the stability field of magmatic epidote early and remained there upon final crystallization. The pluton formed and crystallized at depths greater than ∼30 km. REE data indicate that garnet was absent in the melting region and did not fractionate during crystallization. This suggests that the crust was less than or equal to ∼55 km thick at 188 Ma during the early phases of magmatism in the Coast Plutonic Complex. Late Cretaceous contractional deformation and early Tertiary extension exhumed the rocks to upper crustal levels. Textures of magmatic epidote and other magmatic phases, combined with REE data, can be important for constraining the P–T path followed by magmas.

Introduction

In the Canadian Cordillera, epidote was first recognized as a magmatic phase in the Ecstall pluton, B.C., by Crawford and Hollister (1982). They proposed that the Ecstall pluton must have crystallized at high pressures (>0.6 GPa) on the basis of comparison with experimental work. Zen and Hammarstrom (1984a) reported the presence of magmatic epidote in 14 plutons, in addition to the Ecstall pluton, that occupy a belt along the North American Cordillera (Fig. 1). The presence of magmatic epidote in these plutons was inferred as indicating high-pressure crystallization (>0.8 GPa).

Figure 1.  Map showing the distribution of magmatic epidote-bearing plutons in the North American Cordillera. Modified from Zen and Hammarstrom (1984a), Umhoefer and Schiarizza (1996), and Rusmore et al. (2005).

The stability of magmatic epidote at high pressures is confirmed by experiments with water-saturated tonalite melts that show that at fO₂ conditions buffered by NNO and HM, magmatic epidote is stable at pressures >0.5 GPa and >0.3 GPa respectively (Schmidt and Thompson, 1996; Schmidt and Poli, 2004). The high pressure of crystallization of magmatic epidote can be used to constrain crystallization pressure and therefore tectonic processes (e.g. Zen and Hammarstrom, 1984a).

We describe a newly discovered magmatic epidote-bearing pluton from Mt. Gamsby, within the Coast Mountains (Fig. 1). Although a quartz diorite pluton and a younger tonalite pluton contain textural evidence for magmatic epidote, we only examined the quartz diorite in detail because the tonalite is strongly deformed. We combined textures with thermobarometry and geochemistry to interpret the depth of generation and emplacement of the quartz diorite pluton. We show that textures of magmatic epidote are important for constraining the P–T path of the magma from the source to the intrusion level.

Geological setting

The Canadian Cordillera is composed of terranes accreted to North America during the Mesozoic and Cenozoic (Coney et al., 1980). Two composite terranes, the Intermontane and Insular superterranes (Monger et al., 1982), are separated by the Coast shear zone (CSZ), a near vertical, crustal scale shear zone that parallels much of the length of the Canadian Cordillera and records a polyphase deformation history that includes dextral transpression (85–57 Ma; Andronicos et al., 1999, 2003) and transtension with northeast-side-down normal motion (57–48 Ma; Klepeis et al., 1998; Andronicos et al., 1999) (Fig. 1). The Intermontane and Insular terranes are intruded by plutons of the Coast Plutonic Complex (CPC), which are intermediate in composition (Hutchison, 1982; Samson et al., 1991; Thomas and Sinha, 1999) and formed during subduction of oceanic plates beneath North America (e.g. Engebretson et al., 1985). Magmatic epidote-bearing plutons in the CPC are predominantly late Cretaceous and occur west of the CSZ (Fig. 1; Zen and Hammarstrom, 1984a). The magmatic epidote-bearing plutons at Mt. Gamsby, however, are Jurassic and located east of the CSZ.

Mt. Gamsby is located at a latitude of ∼53^oN (Figs. 1 and 2). The rocks at Mt. Gamsby include three mappable lithologic units: metamorphic country rock and two plutonic bodies (Fig. 2; Hamblock, 2006). Late cross-cutting basalt, basaltic andesite and granodiorite dikes also occur. The country rocks include amphibolite and metasediments with gneissic to migmatitic textures, although schists also occur. Mineral assemblages representative of the rocks at Mt. Gamsby are listed in Table 1. The two plutonic bodies include a 188.1 ± 3.3 Ma quartz diorite and a 155.3 ± 2.7 Ma tonalitic orthogneiss (G.E. Gehrels, personal communication, 2007; Table 2).

Figure 2.  Geological map of Mt. Gamsby in the Coast Mountains of British Columbia. UTM zone 9N and datum WGS 84. Cross-cutting relationships, structural observations and radiometric dating indicate that the amphibolite and metasedimentary rocks are the oldest units in the field area, followed by the quartz diorite pluton and then the tonalite pluton (Hamblock, 2006). The contacts between the country rock and the two plutonic bodies are deformed intrusive contacts. The quartz diorite and tonalite have U/Pb zircon ages of 188.1 ± 3.3 Ma and 155.3 ± 2.7 Ma respectively (G.E. Gehrels, personal communication, 2007) and both have been deformed during a complex deformation history (see Hamblock, 2006). Sample locations for this study are indicated.

Table 1.   Minerals observed in samples from Mt. Gamsby.

Sample	Rock Type	Qtz	Kfs	Pl	Ms	Chl	Bt	Ep	Amph	Grt	Ilm	Mag	Ttn	Rt	Ap	Lm	Hem	Sfd	Clay	Zrn
alt, altered; amph, amphibole; ap, apatite; bt, biotite; chl, chlorite; ep, epidote; grt, garnet; hem, hematite; ilm, ilmenite; kfs, K-feldspar; lm, limonite; mag, magnetite; ms, muscovite; pl, plagioclase; po, pyrrhotite; py, pyrite; qtz, quartz; rt, rutile; ser, sericite; sfd, sulphide; ttn, titanite; zrn, zircon.
05B-4	Amphibole gneiss	×	×	×	×	×	×	×	×			×	×		×					×
05B-9A	Quartz diorite	×	×	×	ser in pl	×	×	×	×		×	×	×		×
05B-11	Tonalite	×		×	×	×	×	×	×		×	×			×					×
05B-13A	Paragneiss	×		×		×	×	×	×		×	×		×	×
05B-13B	Paragneiss	×		×	×	×	× (alt)							×	×		×
05B-13C	Paragneiss	×		×	×	×								×			×
05B-45A	Amphibole gneiss	×	×	×		×	×	×	×		×			×	×	×				×
05B-49	Quartz diorite	×	×	×		×	×	×	×		×				×	×				×
05B-70C	Quartz diorite	×		×	×		×	×	×		×	×	×		×			×
05B-71A	Amphibole gneiss	×	×	×		×	×	×	×	×	×	×				×		po py	×
05B-72	Quartz diorite	×	× (alt)	×		×	×	×	×		×	×	×		×

Table 2.   Major and trace element analyses of Mt. Gamsby rocks.

Sample number	05B-11^1	05B-27^1	05B-30A^2	05B-65^3	05B-70C^3	05B-72^3
*Total Fe reported as FeO. ^1Tonalite; ^2Tonalitic enclave; ^3Quartz diorite.
XRF (wt. %)
SiO2	62.14	62.93	60.58	55.71	53.71	51.30
TiO2	0.48	0.45	0.48	0.99	1.00	1.01
Al2O3	16.92	17.00	17.50	17.54	18.36	18.74
FeO*	5.07	4.65	5.39	8.72	8.45	8.79
MnO	0.14	0.13	0.18	0.16	0.18	0.18
MgO	2.42	2.16	2.33	2.81	3.92	4.61
CaO	5.99	5.61	6.20	8.33	8.19	9.27
Na2O	3.23	3.53	3.47	3.15	3.82	3.62
K2O	1.52	1.56	1.32	0.93	1.06	0.76
P2O5	0.26	0.24	0.29	0.30	0.31	0.32
Total	98.18	98.25	97.75	98.64	99.02	98.62
XRF (ppm)
Ni	10	8	5	12	12	16
Cr	13	11	5	13	16	27
Sc	10	10	11	26	20	23
V	93	79	86	269	213	229
Ba	539	582	527	451	767	580
Rb	55	58	50	19	20	10
Sr	578	590	618	312	731	835
Zr	91	88	87	105	93	57
Y	16	14	21	31	21	19
Nb	4.7	4.9	4.3	1.5	3.6	2.4
Ga	16	15	16	18	19	18
Cu	15	11	20	33	46	22
Zn	54	53	71	88	82	81
Pb	1	2	2	4	3	1
La	11	50	18	10	15	14
Ce	17	96	33	25	32	28
Th	1	10	2	2	1	0
Nd	12	38	17	15	20	18
Modes (vol. %)
Plagioclase	37.1	39.1	33.8		40.5	47.5
Quartz	31.1	30.9	13.1		4.9	8.4
K-feldspar	0.3	2.0	0.0		0.0	0.3
Hornblende	2.1	1.6	4.1		38.8	23.0
Biotite	17.7	21.5	28.8		7.4	10.2
Epidote	11.4	4.9	17.8		7.4	9.0
Chlorite	0.0	0.0	1.6		0.3	0.3
Oxides	0.3	0.0	0.9		0.7	1.5

Description of the plutons

Quartz diorite

The quartz diorite is coarse-grained, salt-and-pepper-coloured and varies in texture from undeformed to mylonitic. It is metaluminous and has SiO₂ contents ranging from 51.30 to 55.71 wt% (Table 2). Polished sections from five samples were examined using petrographic microscope and electron microprobe (samples 05B-9, 05B-9A, 05B-49, 05B-70C, and 05B-72; Fig. 2). Samples 05B-9 and 05B-49 are strongly foliated and are not used to assess crystallization processes. Detailed descriptions of the minerals and the chemistry are presented in subsequent sections.

Tonalitic orthogneiss

In outcrop, the tonalitic orthogneiss is light in colour, varies from relatively unfoliated to mylonitic and contains plagioclase augen and zones of abundant mafic enclaves. Both magmatic and metamorphic epidotes are present. We interpret small, euhedral, inclusions within plagioclase to be magmatic, whereas larger, commonly ragged epidote that occurs oblique to the biotite-defined foliation in the rock appears to be metamorphic. Geochemically, the tonalitic orthogneiss is intermediate in composition, having SiO₂ contents between 62.14 and 62.93 wt% (Table 2).

P–T estimates

Quartz diorite

P–T conditions for crystallization of three quartz diorite samples (05B-9A, 05B-70C, 05B-72) were calculated using hornblende–plagioclase thermometry (Holland and Blundy, 1994) and Al-in-hornblende barometry (Schmidt, 1992). Amphibole is generally pleochroic from straw yellow to blue–green. Some amphibole contains abundant, small, inclusions of magnetite and ilmenite. Amphibole compositions are ferropargasite-pargasite and ferrotchermakite-tchermakite (Table 3; Leake et al., 2004). The composition of the amphibole suggests that it is igneous (Si ≤ 7.5 Leake, 1971; also Ca ≥ 1.6 Hammarstrom and Zen, 1986). Plagioclase is often twinned, oscillatory zoned and shows little evidence of alteration, although a few grains contain sericite. Compositions range from An₂₆ to An₆₁, although most are andesine (Table 3). Core to rim traverses of plagioclase in sample 5B-70C indicate oscillatory zoning with a Na-rich core (oligoclase). Temperature calculations used amphibole–plagioclase rim pairs and pressure calculations used amphibole rim compositions, except for sample 05B-70C in which analyses were averaged to obtain a larger sampling size (Table 3).

Table 3.   *Average hornblende, plagioclase, garnet, and biotite compositions in samples from Mt. Gamsby. Results of thermobarometry calculations using these average values are also shown.

	05B-4		05B-9A		05B-11		05B-13A		05B-13B		05B-13C
	Hbl n = 1	Pl n = 1	Hbl n = 10	Pl n = 5	Hbl n = 9	Pl n = 5	Hbl n = 5	Pl n = 3	Ms n = 6	Pl n = 6	Ms n = 29	Pl n = 19
SiO2	48.98	55.80	40.50	54.77	41.77	60.18	44.19	60.83	45.33	62.06	45.61	62.74
TiO2	0.13	0.00	0.41	0.00	0.51	0.00	0.22	0.00	0.15	0.00	0.30	0.00
Al2O3	7.24	28.19	14.69	29.48	13.82	26.52	12.05	24.33	36.49	24.13	35.57	23.23
FeO	9.49	0.12	21.03	0.16	18.69	0.12	15.41	0.14	1.78	0.06	1.94	0.08
MnO	0.31	0.00	0.40	0.00	0.73	0.01	0.58	0.00	0.02	0.03	0.00	0.00
MgO	17.17	0.01	6.86	0.00	9.07	0.00	11.27	0.07	0.58	0.00	0.62	0.00
CaO	12.52	8.98	11.63	11.43	11.57	6.70	11.11	5.51	0.00	4.66	0.02	4.34
Na2O	0.75	6.59	1.31	5.29	1.27	7.92	1.55	8.42	1.66	8.91	1.45	9.07
K2O	0.14	0.00	0.96	0.05	0.66	0.05	0.24	0.07	9.19	0.09	9.05	0.09
Total	96.73	99.69	97.81	101.19	98.10	101.50	96.63	99.36	95.21	99.96	94.56	99.56
T (°C)	^1714 ± 40 °C		^1797 ± 40 °C		^1735 ± 40 °C		^1688 ± 40 °C		^2**737 °C		^2**716 °C
P (GPa)			^30.94 ± 0.06 °C GPa

	05B-45A		05B-70C		05B-71A			05B-72
	Hbl n = 8	Pl n = 8	Hbl n = 9	Pl n = 4	Grt n = 1	Bt n = 5	Pl n = 9	Hbl n = 17	Pl n = 10
bt, biotite; grt, garnet; hbl, hornblende; ms, muscovite; pl, plagioclase; ^1Holland and Blundy (1994); ^2Green and Usdansky (1986); ^3Schmidt (1992); ^4Ferry and Spear (1978) and Hodges and Spear (1982); ^5Hoisch (1990); *Compositions of minerals were determined using wavelength-dispersive analysis on a Cameca SX50 electron microprobe at the University of Texas-El Paso using a set of natural and synthetic mineral standards, accelerating voltage 15 keV, beam current 20 nA, and beam diameter 2–5 μm; **at 0.8 GPa; 05B-4, 05B-45A; 05B-71A = amphibolite gneiss; 05B-13A, 05B-13B; 05B-13C  =  metasedimentary gneiss; 05B-9A; 05B-70C; 05B-72 = quartz diorite; 05B-11 = tonalite.
SiO2	44.59	59.55	41.84	57.50	37.74	36.04	57.34	42.13	56.42
TiO2	0.46	0.00	0.54	0.03	0.00	1.44	0.00	0.42	0.00
Al2O3	11.53	25.71	13.99	27.74	20.94	17.21	27.67	13.15	27.73
FeO	14.88	0.19	17.71	0.07	26.15	19.34	0.03	16.86	0.11
MnO	0.52	0.00	0.41	0.00	8.73	0.27	0.00	0.34	0.00
MgO	11.50	0.00	9.84	0.00	3.82	10.69	0.00	9.83	0.00
CaO	11.53	6.86	11.57	8.57	4.09	0.05	9.12	11.52	8.97
Na2O	1.32	7.72	1.41	6.75	0.00	0.11	6.55	1.47	6.37
K2O	0.29	0.05	0.72	0.06	0.00	9.11	0.00	0.56	0.03
Total	96.64	100.08	98.04	100.72	101.47	94.27	100.72	96.29	99.64
T (^oC)	^1691 ± 40 °C		^1763 ± 40 °C		^4759–809 ± 50 °C			^1750 ± 40 °C
P (GPa)					^50.96–1.06 ± 0.03 GPa

The thin sections for samples 05B-9A and 05B-72 contain all of the minerals necessary for the appropriate application of the Al-in-hornblende barometer. K-feldspar was not identified in sample 05B-70C. Pressures of 0.94 ± 0.06 GPa and 0.82 ± 0.06 GPa were estimated for samples 05B-9A and 05B-72 respectively. A pressure of 0.86 ± 0.06 GPa was estimated for sample 05B-70C. Temperatures estimated by hornblende–plagioclase thermometry range from 750 to 797 ± 40 °C.

Country rock

A garnet-bearing metasedimentary country rock that occurs as a screen within the quartz diorite was used to estimate pressure by garnet–biotite–plagioclase–quartz barometry (Hoisch, 1990). Garnets in sample 05B-71 show patchy zoning in Mn and Mg and no obvious zoning in Fe. Biotite is unzoned and plagioclase compositions vary between An₃₂ and An₅₄, although most analyses are between An_{43 and} An ₄₅. Average garnet rim and matrix biotite and plagioclase compositions were used for the calculation (Table 3) and a pressure of 0.96–1.06 ± 0.03 GPa was determined (Hamblock, 2006), which is consistent with Al-in-hornblende pressure estimates.

Temperatures estimated using hornblende–plagioclase thermometry (Holland and Blundy, 1994) for the country rock and tonalitic orthogneiss are 688–714 ± 40 °C and 735 ± 40 °C respectively. Temperatures for the country rock estimated using muscovite–plagioclase (Green and Usdansky, 1986; 716–737°C) and garnet–biotite (Ferry and Spear, 1978; Hodges and Spear, 1982; 759–809 ± 50 °C) thermometry are similar to those calculated using hornblende–plagioclase thermometry. The garnet–biotite temperatures are the highest estimated and may be anomalously high owing to possible retrograde re-equilibration and high Mn contents. However, the consistently high temperatures given by several different thermometers in different chemical systems, the presence of migmatites and phase equilibria constraints confirm temperatures in excess of ∼700 °C during pluton emplacement.

Evidence for a magmatic origin of the epidote

Epidote occurs in a variety of textures that we interpret to be magmatic. It occurs as abundant, small (<0.15 mm), highly birefringent, zoned, euhedral to subhedral inclusions in plagioclase (Fig. 3a). Some of these grains appear to be located along cleavage planes, whereas others in the same grain are randomly oriented. We interpret this epidote to be magmatic because the grains have a euhedral habit, the plagioclase is fresh and epidote and plagioclase are zoned. Additionally, the Ps content of epidote inclusions in plagioclase range from 21 to 26 mol%, whereas epidote formed from alteration of plagioclase generally has a lower Ps content (Table 4; Tulloch, 1979: 0–24;  mol%; Carcangiu et al., 1997).

Figure 3.  Photomicrographs showing textures in quartz diorite from Mt. Gamsby. (a) Photomicrograph of sample 05B-72 showing a plagioclase grain containing inclusions of abundant small, euhedral epidote. Photo taken in cross-polarized light (xpl). (b) Photomicrograph of sample 05B-70C showing epidote with a core of allanite. Photo taken in plane-polarized light (ppl). (c) Photomicrograph of sample 05B-70C showing a ragged epidote core surrounded by plagioclase (xpl). The ragged edges suggest early growth and partial resorption in the magma before being enclosed in plagioclase. Epidote with allanite cores that are not included in plagioclase commonly have a rounded habit, which also suggests resorption (e.g., Evans and Vance, 1987). (d) Photomicrograph of sample 05B-9A showing late interstitial biotite relative to amphibole and epidote (ppl). (e) Photomicrograph of sample 05B-9A showing epidote inclusions in biotite (ppl). (f) Photomicrograph of sample 05B-9A showing amphibole with a birefringent core, quartz inclusions and a euhedral rim (ppl). A euhedral apatite is included in the amphibole. Similar textures are observed in magmatic epidote-bearing orbicular diorite from Labrador, where “hornblende commonly contains quartz oikocrysts, suggesting the early crystallization and subsequent resorption of clinopyroxene (Owen, 1991).” Similarly, in New Zealand plutons, “numerous small quartz inclusions occur within hornblende in the most mafic samples, suggesting that the hornblende originally crystallized as clinopyroxene, although no relics have been seen (Allibone and Tulloch, 2004).” In the magmatic epidote-bearing Bushy Point pluton, Alaska, the presence of early crystallizing clinopyroxene was assumed to balance reactions involved in the paragenesis of the pluton, although clinopyroxene was not observed in the rock (Zen and Hammarstrom, 1984b). For all photomicrographs, al = allanite; amph = amphibole; ap = apatite; bt = biotite; ep = epidote; gt = garnet; pl = plagioclase; and qtz = quartz.

Table 4.   Compositions of epidote in quartz diorite from Mt. Gamsby.

Sample	05B-9A	05B-9A	05B-70C	05B-70C	05B-70C	05B-70C	05B-70C	05B-70C	05B-70C	05B-70C	05B-70C	05B-70C	05B-70C	05B-70C	05B-72	05B-72
*All Fe as FeO.
SiO2	38.00	37.63	37.41	37.52	36.84	37.19	37.62	37.60	37.43	37.44	37.66	37.49	36.83	37.43	36.77	37.21
TiO2	0.03	0.02	0.10	0.09	0.03	0.07	0.09	0.14	0.12	0.06	0.05	0.08	0.23	0.12	0.02	0.06
Al2O3	25.34	25.54	25.95	25.62	21.95	24.98	25.92	25.42	25.68	24.93	25.38	25.27	24.51	25.92	24.35	25.10
FeO*	10.09	9.84	10.67	10.75	8.55	11.32	10.29	11.27	11.25	11.42	11.28	11.47	11.92	11.25	11.17	10.32
MnO	0.28	0.34	0.61	0.37	0.27	0.54	0.32	0.19	0.34	0.22	0.36	0.32	0.43	0.38	0.32	0.24
MgO	0.04	0.00	0.00	0.00	0.03	0.00	0.00	0.00	0.01	0.00	0.00	0.00	0.00	0.00	0.00	0.00
CaO	23.47	23.72	23.32	23.64	20.77	23.11	23.48	23.71	23.60	23.45	23.36	23.59	23.23	23.49	22.85	23.54
Na2O	0.00	0.00	0.00	0.00	0.14	0.00	0.04	0.00	0.00	0.00	0.00	0.00	0.00	0.00	0.04	0.01
K2O	0.03	0.04	0.01	0.02	0.03	0.02	0.02	0.03	0.01	0.02	0.02	0.02	0.02	0.02	0.00	0.00
Total	97.28	97.12	98.07	98.01	88.60	97.24	97.79	98.36	98.44	97.54	98.11	98.23	97.17	98.61	95.51	96.48

Epidote also occurs as larger grains (∼1.4 mm) outside or within plagioclase. These grains may be zoned and/or contain allanite cores (Fig. 3b). Some grains within plagioclase have ragged edges (Fig. 3c). Epidote with allanite cores that are not included in plagioclase commonly has a rounded habit. Although both zoning and allanite cores are commonly interpreted as indicating a magmatic origin of epidote (Zen and Hammarstrom, 1984a; Schmidt and Poli, 2004), secondary epidote may grow around allanite and thus an allanite core alone does not indicate a magmatic origin (e.g., Liu et al., 1999). Some epidote in the quartz diorite contains both zoning and allanite cores, which strongly suggests a magmatic origin. Furthermore, secondary Ca-silicate minerals are not observed in the quartz diorite, as they are in rocks containing secondary epidote rimming allanite (Liu et al., 1999).

Epidote in the quartz diorite also occurs as euhedral to subhedral grains within biotite (Fig. 3d, e), a texture commonly interpreted as magmatic (e.g. Zen and Hammarstrom, 1984a). This type of epidote in the quartz diorite is dissimilar to secondary epidote after biotite, which is anhedral, elongate along cleavage planes, associated with other secondary minerals, and has high Ps contents (e.g. Tulloch, 1979; : Ps 36–48 mol%; Dawes and Evans, 1991; : Ps 29–33 mol%; Leake, 1998; Allibone and Tulloch, 2004).

Crystallization sequence and interpretation

Here we present information on the crystallization sequence in the Mt. Gamsby quartz diorite. We compare this information with experimentally determined stability fields for magmatic epidote and other magmatic minerals in tonalite melts (Schmidt and Thompson, 1996; Schmidt and Poli, 2004). The phase equilibria for tonalite melts have been used because those for quartz diorite melts have not yet been published.

The stability field for garnet in water-saturated tonalite melts is greater than ∼1.4 GPa (Fig. 4). Chondrite-normalized rare-earth element (REE) diagrams for the quartz diorite do not show steep negative slopes [(La/Yb)_N = 2.13–5.74], as would be expected if garnet had been present as a restite phase or fractionated from the melt (Fig. 5). Although garnet could have been present and subsequently resorbed by the melt, this seems unlikely. Garnet is a dense phase and thus has a tendency to accumulate; additionally, at high pressures, garnet is stable as a restite phase (Rapp and Watson, 1995). Therefore, the lack of heavy REE depletion in the quartz diorite suggests that melting occurred at pressures below the stability of garnet.

Figure 4.  Pressure-temperature graph showing experimental phase equilibria for water-saturated tonalite liquid (Schmidt and Thompson, 1996; Schmidt and Poli, 2004). Stippled region shows the stability field for magmatic epidote. The P–T point for sample 05B-71A is from garnet–biotite thermometry (Ferry and Spear, 1978; Hodges and Spear, 1982) and garnet–biotite–plagioclase–quartz barometry (Hoisch, 1990) (Hamblock, 2006). The P–T point for sample 05B-9A is from hornblende–plagioclase thermometry (Holland and Blundy, 1994) and Al-in-hornblende barometry (Schmidt, 1992). The aluminosilicate triple point of Pattison (1992) is shown for reference. The arrow shows the P–T path for the quartz diorite deduced from mineral textures in the pluton. The circles containing letters refer to photomicrographs in Fig. 3 that contain textures representative of the indicated P–T conditions.

Figure 5.  Chondrite-normalized REE diagram for quartz diorite samples from Mt. Gamsby. Normalization values are from McDonough and Sun (1995).

Whereas garnet provides constraints on pressure, the presence of clinopyroxene provides constraints on temperatures. In tonalitic magmas, clinopyroxene is stable at temperatures >770 °C (Fig. 4; Schmidt and Thompson, 1996; Schmidt and Poli, 2004). Although clinopyroxene is not observed in our samples, the textures of some of the amphiboles suggest that they formed through incongruent hydration reactions involving clinopyroxene. Some amphiboles have cores with higher order birefringence than their rims. These cores also contain quartz inclusions (Fig. 3f). These amphiboles are similar to those with quartz inclusions from the Bell Island tonalite, Alaska, which were interpreted as having formed by hydration crystallization by the reaction:

• (1)

Others have inferred the presence of early magmatic clinopyroxene in magma despite the absence of the clinopyroxene in the rock on the basis of similar amphibole textures (Zen and Hammarstrom, 1984b; Owen, 1991; Allibone and Tulloch, 2004). We interpret the amphibole with highly birefringent cores and quartz inclusions as a reaction product of pyroxene and melt, which indicates the presence of early magmatic pyroxene.

In the quartz diorite, epidote occurs as inclusions in plagioclase and biotite. Crystallization of epidote before plagioclase occurs at pressures greater than ∼1 GPa, and crystallization of epidote before biotite occurs at pressures greater than ∼0.9 GPa (Fig. 4; Schmidt and Poli, 2004). Biotite appears to be late in the crystallization sequence, as it commonly occurs as anhedral grains that surround amphibole and contains inclusions of epidote (Fig. 3d, e). Biotite shows a pleochroism from straw yellow to dark brown and is generally unaltered.

Taken together, textures and chemical data indicate that crystallization occurred at pressures less than ∼1.4 GPa and cooled from temperatures >770 °C nearly isobarically to the solidus (Fig. 4). A final crystallization pressure of ∼0.87 GPa based on Al-in-hornblende barometry confirms pressures estimated from epidote textures.

Discussion

Using magmatic epidote to deduce P–T path

Schmidt and Poli (2004, p. 420) showed the potential of using the crystallization sequence in magmatic epidote-bearing rocks as a barometer. The lack of garnet and evidence for clinopyroxene as an early crystallizing phase in samples of quartz diorite from Mt. Gamsby indicates that the P–T of the source region was less than ∼1.4 GPa and greater than 770 °C (e.g. Schmidt and Poli, 2004). Other textural relationships in the rock indicate that the magma then passed through the epidote-in, plagioclase-in, and biotite-in fields respectively, which follows a path of decreasing temperature at decreasing or constant pressure (Fig. 4). Al-in-hornblende barometry indicates that the final crystallization of the quartz diorite occurred at pressures of ∼0.87 GPa.

Tectonic implications

The emplacement depth for the 188.1 ± 3.3 Ma quartz diorite is ∼30 km, based on epidote textures and Al-in-hornblende estimates of ∼0.87 GPa. The lack of evidence for garnet in the quartz diorite indicates that the plutons were generated at pressures <1.4 GPa. If the magmas at Mt. Gamsby were sourced near the base of the crust, then the crust at 188 Ma was less than ∼55 km thick and the plutons were emplaced at middle to lower crustal depths.

Magmatic epidote in the 155.3 ± 2.7 Ma tonalite orthogneiss at Mt. Gamsby suggests that the rocks had not been substantially exhumed by this time. The magmatic epidote-bearing plutons in the Coast Mountains thus indicate that the crust was less than or equal to ∼55 km thick at 188 Ma during the early phases of magmatism and probably did not substantially thicken until after 155 Ma, after emplacement of the tonalite orthogneiss at Mt. Gamsby. The emplacement of voluminous magmatic epidote-bearing plutons between 110 and ∼90 Ma suggests that similar crustal thickness may have persisted throughout much of the magmatic history of the CPC.

Importantly, after ∼90 Ma, magmatic epidote-bearing plutons are absent, suggesting that conditions within the arc changed. This is consistent with the vigorous tectonic activity that occurred between 90 and 48 Ma, as is established in the published literature on the tectonic evolution of this segment of the Canadian Cordillera, and includes: (1) long-lived magmatism within an arc or arcs between ∼180 and 90 Ma (e.g. Chardon et al.,1999; and references therein); (2) major transpression, thrusting and crustal thickening coupled with uplift and exhumation between 90 and 60 Ma (Hollister and Crawford, 1986; Crawford et al., 1987; Andronicos et al., 1999) and (3) crustal extension and exhumation during detachment faulting between 60 and 48 Ma (Hollister and Andronicos, 2000; Andronicos et al., 2003).

Our study suggests that the CPC represents a stable, long-lived arc or arcs in which intermittent magmatism occurred between ∼188 and 90 Ma, including intrusion of magmatic epidote-bearing plutons that cooled nearly isobarically at depth. The crust during this time does not appear to be hugely over thickened, but relatively thin for orogenic belts, and did not over-thicken until after the emplacement of the magmatic epidote-bearing plutons.

Acknowledgements

Minghua Ren and Gabriela Depine provided excellent assistance in acquiring microprobe data and José Hurtado, Jr., provided assistance in the field. We thank James Beard, Lincoln Hollister and two anonymous reviewers for thoughtful reviews of an earlier draft. We also thank Gerhard Franz and two anonymous reviewers for helpful reviews of the current draft. This work was funded by NSF Grant No. EAR-0310347 to Andronicos.