JOURNAL OF PETROLOGY
VOLUME 38
NUMBER 11
PAGES 1489–1512
1997
Fluid Flow and Diffusion in the Waterville
Limestone, South–Central Maine:
Constraints from Strontium, Oxygen and
Carbon Isotope Profiles
M. J. BICKLE1∗, H. J. CHAPMAN1, J. M. FERRY2, D. RUMBLE III3 AND
A. E. FALLICK4
1
DEPARTMENT OF EARTH SCIENCES, UNIVERSITY OF CAMBRIDGE, DOWNING STREET, CAMBRIDGE CB2 3EQ, UK
2
DEPARTMENT OF EARTH AND PLANETARY SCIENCES, JOHNS HOPKINS UNIVERSITY, BALTIMORE, MD 21218, USA
3
GEOPHYSICAL LABORATORY, CARNEGIE INSTITUTION, WASHINGTON, DC 20015, USA
4
SURRC, EAST KILBRIDE, GLASGOW G75 0QF, UK
RECEIVED DECEMBER 12, 1996; REVISED TYPESCRIPT ACCEPTED JUNE 20, 1997
Oxygen, carbon and strontium isotopic profiles across the margin of
the Waterville limestone member are used to investigate advective
and diffusive transport during metamorphism of the Waterville
Formation in south–central Maine, USA. Rb–Sr isotopic systematics were homogenized on the ~10 cm hand-specimen scale at
ages that are within error of the 376±6 Ma Rb–Sr whole-rock
age of the syn-metamorphic Hallowell pluton. This is consistent
with a plutonic heat source for this low-pressure andalusite- and
sillimanite-grade Acadian metamorphic terrane. Advective displacements of all three isotope profiles at the garnet-grade Blue Rock
Quarry indicate fluid flow to the east into the limestone, and the
oxygen-isotope profile implies a time-integrated fluid flux of
3·2±1·4 m3/m2 (2r error). This cross layer flux is insufficient
to cause the observed reaction progress of the muscovite+
ankerite+quartz to biotite+anorthite (in plagioclase)+calcite reaction in the ~100 m thick Waterville limestone member and much
of the fluid flow responsible may have been layer parallel. The
isotope profiles indicate advective–diffusive homogenization over
distances of 1·5 m (d13C) to 6 m (d18O) and such homogenization
distances are difficult to reconcile with observations of order of
magnitude variations in reaction progress on the centimetre scale or
less. It is possible that infiltration occurred during events short lived
compared with diffusion, that the reactions started at different
temperatures dependent on bulk composition or that diffusion of
water from layers with less reactants to layers with more reactants
was important in driving the biotite-producing reaction. However,
∗Corresponding author. Fax: 01223 333450. e-mail: [email protected].
ac.uk
variations of fluid composition inferred from the mineral assemblages
are apparently inconsistent with diffusion driving reaction progress,
and models of precursor assemblages do not indicate significant
compositional control of the temperature of the first appearance of
biotite in the rocks. Irrespective of the details of flow and diffusive
exchange on the centimetre scale, the average reaction progress in the
Waterville limestone member requires significant layer-parallel fluid
fluxes.
KEY WORDS: metamorphism; fluid flow; strontium isotopes; oxygen isotopes;
carbon isotopes
INTRODUCTION
Regional metamorphism of pelitic rocks is associated
with loss of ~10% by volume of an H2O–CO2 fluid
phase (Walther & Orville, 1982). The resultant fluid flux,
possibly augmented by external fluid inputs, may control
the chemical, petrological, thermal and deformational
evolution of metamorphic crust (e.g. Garlick & Epstein,
1967; Rye et al., 1976; Etheridge et al., 1983; Hoisch, 1987;
Ague, 1994; Ferry, 1994) as well as causing significant
 Oxford University Press 1997
JOURNAL OF PETROLOGY
VOLUME 38
geochemical exchange between the solid Earth and the
atmosphere and hydrosphere (Berner et al., 1983; Kerrick
& Caldeira, 1993; Bickle, 1994, 1996). Despite its importance, the magnitude of metamorphic fluid fluxes and
fluid transport mechanisms through metamorphic rocks
are controversial [compare Ferry (1994) and Goodge &
Holdaway (1995)].
A number of chemical, isotopic and petrological indices
of fluid movement have been used to investigate fluid
movement in metamorphic rocks. Measurement of isotopic profiles across boundary layers in strata of contrasting compositions have been successful in monitoring
a component of the time-integrated fluid flux and constraining diffusive exchange, time, porosity and fluid–rock
exchange mechanisms (e.g. Rye et al., 1976; Ganor et al.,
1989; Bickle & Baker, 1990a; Bickle, 1992). However,
the majority of the fluid flux in metamorphic rocks is
probably layer parallel, being channeled as a consequence
of the large permeability contrasts between layers (e.g.
Bickle & Baker, 1990a; Skelton et al., 1995; Yardley
& Lloyd, 1995). Pinned-boundary type boundary-layer
isotopic profiles, in which the isotopic composition of the
high-flux layer remains uniform right up to the contact,
reveal this but cannot resolve the layer-parallel component of the time-integrated flux. The index that should
record the total fluid-flow history of a rock is the reaction
progress of a mineral assemblage whose stability is sensitive to fluid composition. Ferry (e.g. 1992, 1994) has
made much use of reaction progress in impure carbonate
rocks infiltrated by H2O–CO2 fluids to investigate fluid
flow in Acadian metamorphic terranes in New England.
He found evidence for large time-integrated fluid fluxes
(102−104 m3/m2) in several terranes, concluded that flow
direction was near horizontal and up temperature, and
found evidence for layer-parallel flow with order-ofmagnitude variations in time-integrated flux in adjacent
centimetre-thick layers. These conclusions are controversial. Wood & Graham (1986) questioned the sensitivity of the calculations to the accuracy of the
geothermometry. The modelling presumes that mineral
assemblages grow in near chemical equilibrium with the
fluid phase on the grain scale. If the fluid–solid reactions
were kinetically limited the interpretation might be very
different (e.g. Lasaga & Rye, 1993). Metamorphic rock
will compact and expel fluid upwards relatively rapidly
(Skelton et al., 1997b). The mechanism which could
maintain near-horizontal up-temperature flow over distances of tens of kilometres is not known.
In this paper we attempt to test the reaction-progress
based interpretations of fluid flow regimes in south–
central Maine (e.g. Ferry, 1987, 1994; Baumgartner &
Ferry, 1991) by examination of strontium, oxygen and
carbon isotope profiles across the margin of a ~100 m
thick limestone unit (Waterville limestone member,
Fig. 1). The isotope measurements place constraints on
NUMBER 11
NOVEMBER 1997
Fig. 1. Geological map and metamorphic isograds in Waterville and
Sangerville Formations, south–central Maine, USA, after Ferry (1988).
Arrows show time-integrated fluid flux estimates of Baumgartner &
Ferry (1991). Ticks on isograds denote side of mineral appearance.
Biotite, garnet, staurolite, andalusite and sillimanite isograds are in
pelitic rocks, and amphibole, zoisite and diopside isograds are in calcsilicate rocks.
the timing of fluid flow, magnitude of cross-layer flow
and the kinetics of fluid–solid exchange. The results are
compared with estimates of reaction progress and fluid
composition from the mineral assemblages.
GEOLOGICAL SETTING
The meta-sedimentary Sangerville and Waterville Formations of south–central Maine were isoclinally folded,
intruded by peraluminous granitic plutons and metamorphosed to grades between chlorite and sillimanite
zones during the Devonian Acadian orogeny (Osberg,
1968, 1988; Fig. 1). The metamorphic fabric overgrows
the dominant steeply dipping, north-east striking foliation
which is axial planar to the tight to isoclinal folds. The
Sangerville Formation is predominantly a calcareous
1490
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
greywacke and the Waterville Formation is mostly pelitic
schist with minor micaceous sandstone and rare 1–5 cm
limestone beds. One ~100 m thick limestone unit (informally called the Waterville limestone member) can be
traced the length of the formation and outlines major
isoclinal folds (Fig. 1). This is a thinly bedded carbonaterich marl. The contact between the Waterville limestone
and the adjacent pelitic rocks is the focus of much of this
paper. The Sangerville and Waterville Formations are
thought to be Silurian, on the basis of sparse fossil
evidence (Osberg, 1968, 1988). Three of the granitic
plutons (Hallowell, Togus and Three Mile Pond plutons)
have been dated by whole-rock Rb–Sr isochrons at
387±11, 394±8 and 381±14 Ma (Dallmeyer & Van
Breeman, 1981). The low-pressure andalusite–sillimanite
facies series metamorphism increases in grade from northeast to south-west with the highest metamorphic grade
assemblages (diopside in calc-silicates, sillimanite in pelites) developed around the granite plutons. Ferry (1976,
1980) mapped metamorphic isograds by the appearance
of index minerals and concluded from the geometrical
relationship between the regional metamorphic isograds
and the plutons that the metamorphism was caused by
heat advected by granite intrusion.
SAMPLING AND ANALYTICAL
METHODS
Sample localities are shown in Fig. 1. Detailed profiles
across Waterville limestone and pelite perpendicular to
the contact were sampled at localities 5 (garnet grade)
and 7 (chlorite grade). Whole-rock samples weighing
between 500 g and 2 kg were collected at spacings of
~5 cm adjacent to contacts and at progressively wider
spacings further from contacts (Table 1). Sample spacings
are given from the contact between phyllite and limestone
for locality 7. The contact at locality 5 was located on
the basis of whole-rock analyses of samples at –20 cm
with respect to the measured datum for the profile at
locality 5. This locality shows continuous gradation from
a carbonate marl to phyllite over a distance of ~0·5 m,
as discussed below. Some samples were sawn into ~1 cm
slices, which were crushed and analysed separately. In
addition, samples across thin carbonate bands in the
Sangerville Formation were collected at localities 117
and 389 and from Waterville Formation at outcrops 3,
411 and 674. Whole-rock samples were sawn to remove
weathered material and washed in distilled water before
crushing in a soft-iron jaw crusher followed by fine
crushing of an ~80 g aliquot in an agate swing mill.
Rb, Sr and 87Sr/86Sr data are listed in Table 1. Sr
isotopic analyses were performed on the VG54E mass
spectrometer at Cambridge and analyses of NBS
SRM987 standard gave a mean of 0·710253±50 (2r)
over the period of the analyses. Analytical, chemical
processing methods and isotopic spikes were the same as
those used by Bickle et al. (1988). Sr blanks were <1 ng and
are negligible. Rb/Sr analyses were by X-ray fluorescence
(XRF) on pressed powder pellets carried out by P. Webb
at the Open University calibrated against USGS standards [values of de Laeter & Abercrombie (1970)]. Rb/
Sr ratios reproduce to within 2% of the ratio or, for low
Rb samples, with Rb concentrations to no better than
±0·4 p.p.m. of standard values. Samples for isotopic
analysis were ignited in a furnace at 900°C for 8 h, and
treated with HF (+HNO3), HNO3 and HCl in bombs
at 200°C for 8 h each step. Despite this treatment, some
dark material of presumed organic origin failed to either
ignite or dissolve even when treated with HClO4 or after
repeated stages of HNO3. Repeat analyses with this
sample treatment (but not with any of the steps omitted)
gave consistent results and it is presumed that solid and
reagents reached strontium isotopic equilibrium during
the dissolution treatment. Errors are quoted at 2r except
where stated. Isochron regressions are calculated after
York (1969).
Oxygen and carbon isotopic analyses of calcites
(Table 1) were performed at the Geophysical Laboratory,
Washington, in collaboration with D. Rumble. Analytical
methods were as described by Rumble et al. (1991) and
involved reaction with 100% phosphoric acid at 25°C
for time periods between 10 min and 16 h in two-legged
reaction vessels. Values are reported relative to VSMOW
(d18O) and VPDB (d13C). Results were calibrated against
NBS-18 (d18O=7·2, d13C=–0·50) and NBS-19 (d18O=
28·65, d13C= +1·92). Waterville limestone contains significant ankerite (Ferry, 1987) and longer reaction times
may have decomposed ankerite in addition to calcite.
However, sample duplicates and measurements on centimetre-sliced carbonate samples ( JF5-79, -80, -91 and
-92, Table 1) showed good reproducibility, with the mean
1r deviation of the centimetre-sliced samples being
0·27‰. Quartz mineral separates were purified in
HFSiO4 and their oxygen isotope compositions measured
by A. F. Fallick at SURRC using the fluorination method
of Clayton & Mayeda (1963) as modified for ClF3 (Borthwick & Harmon, 1982). Sample size was 10 mg and
reaction temperature 650°C for 15 h. Reaction yields
were checked manometrically on cryogenically purified
CO2. Isotope ratios were determined with a VG Isogas
SIRA 10 triple-collector mass spectrometer. All quartz
data are reported relative to VSMOW; analytical precision (1r) is estimated as ±0·2‰ or better and NBS28 gives d18O=9·6.
Mineral compositions were analysed using the JEOL
JXA-8600 electron microprobe at Johns Hopkins University by J. M. Ferry using natural mineral standards
and a ZAF data correction scheme (Armstrong, 1989).
Major element whole-rock analyses were obtained by
1491
JOURNAL OF PETROLOGY
VOLUME 38
NUMBER 11
NOVEMBER 1997
Table 1: Rb, Sr, O and C isotopic analyses
Sample
Rb1
Sr1
Rb/Sr1
87
Rb/86Sr
87
Sr/86Sr
1r
Distance2 (m) d18O3
d13C
Rock type5
Locality 7, Waterville Bridge
JF7A-1a
19·7
808·1
0·0244
0·0705
0·710218
15
0·01
18·2
−1·9
C
JF7A-1b
20·4
795·7
0·0256
0·0742
0·710190
12
0·02
18·8
−1·9
C
JF7A-1c
15·0
867·8
0·0173
0·0502
0·709992
15
0·03
18·3
−1·9
C
JF7A-2
42·4
904·4
0·0470
0·1360
0·710404
17
0·06
18·2
−2·0
C
JF7A-2a
57·9
868·5
0·0667
0·1930
0·710376
12
0·06
18·4
−1·5
C
JF7A-2av
51·0
1186·8
0·0430
0·1244
0·710152
11
0·06
18·3
−1·5
C
JF7A-3a
87·5
707·5
0·1236
0·3578
0·711287
11
0·14
18·6
−1·6
C
JF7A-3b
94·3
698·9
0·1350
0·3907
0·711464
10
0·14
18·6
−1·7
C
JF7A-3c
66·1
723·3
0·0913
0·2643
0·710786
12
0·14
18·4
−1·7
C
JF7A-3d
33·5
902·0
0·0372
0·1076
0·710066
16
0·14
18·3
−1·7
C
JF7A-3e
52·8
774·0
0·0682
0·1975
0·710474
17
0·14
18·5
−1·7
C
JF7A-3f
14·8
985·3
0·0151
0·0436
0·709732
12
0·14
18·3
−1·8
C
JF7A-4
96·5
734·7
0·1313
0·3800
0·711423
11
0·19
18·8
−1·7
C
JF7A-5
78·4
633·0
0·1238
0·3584
0·711328
13
0·26
18·8
−1·7
C
JF7A-6
48·7
1077·6
0·0452
0·1307
0·710017
14
0·43
18·4
−1·6
C
JF7A-7
33·7
1048·2
0·0322
0·0931
0·709508
29
0·72
18·6
−0·9
C
JF7A-8
23·3
1194·8
0·0195
0·0565
0·709268
10
0·91
18·6
−0·8
C
JF7A-8v
21·9
1207·7
0·0181
0·0525
0·709205
17
0·91
18·7
−0·8
C
JF7A-9
62·3
683·4
0·0912
0·2639
0·710446
12
1·56
18·9
−0·8
C
JF7A-10
19·2
1201·5
0·0160
0·0462
0·709032
23
2·64
18·9
−0·5
C
JF7A-12
17·2
1240·1
0·0138
0·0401
0·709135
9
4·83
19·0
−0·5
C
JF7A-13a
14·6
1184·0
0·0123
0·0356
0·709132
10
6·63
18·9
−0·5
C
JF7A-13b
5·7
1242·2
0·0046
0·0133
0·708852
16
6·63
19·0
−0·5
C
JF7A-13c
6·7
1235·3
0·0054
0·0156
0·708891
13
6·63
18·8
−0·6
C
JF7A-13d
11·4
1133·5
0·0100
0·0289
0·708984
14
6·63
19·0
−0·5
C
JF7A-13e
9·2
1194·6
0·0077
0·0223
0·708953
14
6·63
19·0
−0·5
C
JF7A-13f
8·3
1167·3
0·0071
0·0205
0·708940
10
6·63
19·0
−0·5
C
JF7a-14
33·8
531·1
0·0636
0·1841
0·709863
24
6·63
19·0
−0·6
C
JF7A-15a
21·4
1295·4
0·0165
0·0478
0·708982
11
11·96
19·0
−0·6
C
JF7A-15b
29·2
1303·7
0·0224
0·0648
0·709108
13
11·96
19·1
−0·7
C
JF7A-15c
23·1
1284·3
0·0180
0·0521
0·709139
18
11·96
18·8
−0·6
C
JF7A-15d
21·5
1229·5
0·0175
0·0506
0·709045
23
11·96
18·9
−0·6
C
JF7A-16
22·2
1106·2
0·0200
0·0579
0·709300
15
15·16
18·9
−0·6
C
JF7A-18
34·0
1210·9
0·0281
0·0813
0·708889
14
17·30
19·0
−0·6
C
JF7A-19
99·1
452·6
0·2189
0·6336
0·711430
17
17·91
19·4
−0·7
C
JF7A-21
7·7
1302·0
0·0059
0·0171
0·708837
17
22·05
18·9
−0·5
C
JF7A-22
100·4
97·4
0·0307
2·9878
0·726775
9
29·44
19·6
−3·4
C
JF7A-23
57·3
179·1
0·3197
0·9258
0·716256
17
39·62
18·5
−4·0
C
JF7A-24a
21·5
1201·4
0·0179
0·0518
0·708733
13
52·43
19·4
−0·3
C
JF7A-24c
22·7
1137·6
0·0199
0·0576
0·708824
12
52·43
—
—
C
G
JF7A-25
15·8
916·8
0·0172
0·0498
0·708830
18
64·01
19·5
−0·7
JF7A-26v
14·9
1248·4
0·0119
0·0344
0·708712
12
64·01
19·5
−0·7
C
JF7A-27a
14·6
1272·6
0·0115
0·0333
0·708741
22
79·55
—
—
C
JF7A-27b
18·4
1218·7
0·0151
0·0436
0·708812
21
79·55
—
—
C
JF7A-27c
18·7
1219·9
0·0153
0·0443
0·708783
22
79·55
19·8
−0·5
C
JF7A-27d
13·8
1270·2
0·0109
0·0315
0·708661
12
79·55
—
—
C
JF7A-27e
12·2
1246·0
0·0098
0·0284
0·708687
12
79·55
—
—
C
1492
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
Rb/Sr1
87
87
Sr/86Sr
1r
Distance2 (m) d18O3
d13C
Rock type5
518·0
0·2000
0·5790
0·712986
28
−0·01
—
—
P
448·0
0·1560
0·4516
0·712421
12
−0·10
—
—
P
110·6
363·8
0·3041
0·8805
0·714789
13
−0·15
—
—
P
JF7A-31
117·6
303·0
0·3881
1·1238
0·716279
13
−0·27
—
—
P
JF7A-32
97·7
216·0
0·4522
1·3096
0·717573
24
−0·43
—
—
P
JF7A-33a
82·0
110·7
0·7411
2·1473
0·722319
18
−0·56
—
—
P
JF7A-33b
89·6
174·0
0·5148
1·4911
0·718783
11
−0·56
—
—
P
JF7A-33c
81·2
101·3
0·8013
2·3221
0·723693
18
−0·56
—
—
P
JF7A-33d
95·6
156·5
0·6108
1·7694
0·720230
22
−0·56
—
—
P
JF7A-33e
83·0
94·3
0·8803
2·5512
0·724657
22
−0·56
—
—
P
JF7A-33f
93·5
112·0
0·8351
2·4200
0·723514
24
−0·56
—
—
P
JF7A-34
72·9
65·0
1·1215
3·2516
0·728796
12
−0·97
—
—
P
JF7A-35a
92·9
348·8
0·2664
0·7713
0·714495
15
−1·30
—
—
P
JF7A-35b
111·6
444·6
0·2509
0·7264
0·714185
15
−1·30
—
—
P
JF7A-35c
131·9
257·0
0·5133
1·4867
0·718247
12
−1·30
—
—
P
JF7A-35d
124·4
271·4
0·4583
1·3273
0·717383
17
−1·30
—
—
P
JF7A-35e
142·2
291·9
0·4871
1·4108
0·717948
12
−1·30
—
—
P
JF7A-35f
108·8
333·2
0·3267
0·9460
0·715385
18
−1·30
—
—
P
JF7A-364
88·0
353·6
0·2488
0·7204
0·714680
14
−0·01
—
—
P
JF7A-374
123·0
153·2
0·8484
2·4586
0·724124
25
−0·07
—
—
P
Sample
Rb1
JF7A-28
103·6
JF7A-29
69·9
JF7A-30
Sr1
Rb/86Sr
JF7A-384
92·9
326·8
0·2844
0·8235
0·715324
20
−0·11
—
—
P
JF7A-394
133·6
181·3
0·7368
2·1348
0·722251
15
−0·23
—
—
P
Locality 5, Blue Rock Quarry
−6·4
0·75
−6·3
0·02
—
—
0·85
18·9
−4·6
0·87
−6·4
0·07
JF5-79a
140·2
437·6
0·3204
0·9277
0·716363
12
0·10
17·6
JF5-79b
144·9
415·0
0·3492
1·0114
0·716827
21
0·12
17·7
JF5-79c
156·4
465·3
0·3362
0·9737
0·716761
25
0·16
JF5-80a
150·2
394·9
0·3803
1·1014
0·717626
11
0·01
JF5-80b
141·4
401·3
0·3524
1·0204
0·717064
16
0·02
17·6
JF5-80c
141·1
395·5
0·3568
1·0332
0·717247
12
0·07
18·2
−5·6
0·82
JF5-81a
171·0
351·0
0·4873
1·4116
0·719804
15
0·0
—
—
0·91
JF5-81b
176·7
315·4
0·5604
1·6235
0·720911
21
−0·10
—
JF5-82a
134·2
324·3
0·4138
1·1985
0·718330
24
0·26
—
JF5-82b
111·8
366·1
0·3054
0·8844
0·716635
21
0·26
—
JF5-83
140·3
651·2
0·2154
0·6237
0·714623
14
−0·17
17·6
JF5-84a
87·4
929·6
0·0940
0·2720
0·712102
26
−0·48
JF5-84b
74·9
1044·2
0·0717
0·2075
0·711779
15
JF5-86
66·9
838·6
0·0797
0·2308
0·711039
45
JF5-87
108·6
601·0
0·1807
0·5231
0·712768
19·23q
18·61q
18·91q
—
0·92
—
0·96
18·58q
—
0·87
17·74q
−4·2
0·72
17·3
−4·4
0·40
−0·48
—
—
0·30
−0·97
17·9
−2·5
0·25
12
−1·83
17·9
−3·2
0·50
0·09
JF5-88
39·0
758·2
0·0514
0·1488
0·709987
13
−2·49
18·5
−0·8
JF5-89
28·5
741·9
0·0384
0·1111
0·709648
10
−4·22
18·8
−0·6
0·03
JF5-90
22·8
725·4
0·0315
0·0912
0·709331
20
−6·35
19·2
−0·1
0·00
JF5-91a
25·3
867·3
0·0292
0·0845
0·709256
17
−9·32
19·3
−0·3
0·00
JF5-91b
19·3
914·9
0·0211
0·0611
0·709046
16
−9·32
19·3
−0·2
0·00
JF5-91c
139·7
710·3
0·1967
0·5693
0·711771
16
−9·32
19·4
−0·1
0·00
JF5-91d
19·2
835·0
0·0229
0·0663
0·709120
19
−9·32
19·7
−0·7
0·47
1493
JOURNAL OF PETROLOGY
VOLUME 38
NUMBER 11
NOVEMBER 1997
Table 1: continued
Sample
Rb1
Sr1
Rb/Sr1
87
Rb/86Sr
87
Sr/86Sr
1r
Distance2 (m) d18O3
d13C
Rock type6
0·18
JF5-91e
74·4
892·1
0·0834
0·2414
0·710102
14
−9·32
19·4
−0·4
JF5-91g
55·8
754·5
0·0740
0·2142
0·709894
17
−9·32
19·4
−0·4
0·00
JF5-92a
15·8
794·4
0·0198
0·0574
0·709142
13
−14·63
18·6
−0·1
C
JF5-92b
26·5
759·3
0·0349
0·1009
0·709391
12
−14·63
18·8
−0·2
C
JF5-92c
67·8
529·9
0·1279
0·3702
0·710880
16
−14·63
19·0
−0·7
0·22
JF5-92d
41·2
558·6
0·0738
0·2135
0·710024
18
JF5-92e
−14·63
19·0
−0·6
C
−14·63
19·0
−0·5
C
JF5-93
20·4
978·5
0·0209
0·0605
0·708834
20
−26·82
19·8
−0·4
0·00
JF5-94
54·0
737·6
0·0731
0·2115
0·709913
12
−40·06
19·9
+0·1
0·15
JF5-95
65·6
650·7
0·1008
0·2917
0·710332
24
−56·52
19·8
−0·6
0·18
JF5-96
45·4
849·1
0·0534
0·1545
0·709470
14
−76·20
20·1
−0·7
0·13
JF5-97
197·3
355·5
0·5551
1·6082
0·721011
12
0·14
—
16·72q
—
0·96
JF5-98
225·0
423·0
0·5320
1·5413
0·721130
13
0·30
—
16·80q
—
0·97
JF5-99
244·3
262·8
0·9295
2·6943
0·726645
15
0·58
—
18·92q
—
1·00
JF5-100
158·2
189·6
0·8343
2·4181
0·725346
19
0·86
—
17·66q
—
1·00
JF5-101
131·0
276·8
0·4731
1·3705
0·720213
20
1·47
—
19·00q
—
0·96
JF5-102
165·8
206·6
0·8026
2·3261
0·725028
14
2·26
—
19·08q
—
1·00
JF5-103
270·9
170·5
0·5889
4·6101
0·736265
23
3·20
—
—
1·00
JF5-104a
145·0
326·1
0·4446
1·2878
0·718686
12
4·47
17·1
−6·6
0·78
JF5-104b
141·9
333·6
0·4253
1·2318
0·718459
12
4·47
16·9
−6·6
0·78
JF5-104c
140·6
332·3
0·4233
1·2260
0·718348
15
4·47
17·2
−6·5
0·79
JF5-104d
145·6
347·1
0·4193
1·2144
0·718429
15
4·47
17·6
−6·7
0·87
JF5-104e
142·7
332·4
0·4291
1·2428
0·718498
31
4·47
17·3
−6·6
0·78
JF5-105
151·7
134·5
1·1279
3·2711
0·731927
11
6·30
—
18·25q
—
1·00
JF5-106a
151·5
301·5
0·5025
1·4556
0·719996
14
10·92
—
—
—
0·95
JF5-106b
152·1
284·7
0·5344
1·5481
0·720532
27
10·92
—
18·56q
—
P
JF5-106c
143·8
301·8
0·4763
1·3797
0·719679
15
10·92
—
—
P
JF5-106d
152·2
291·2
0·5228
1·5145
0·720335
11
10·92
—
—
P
JF5-106e
156·6
263·0
0·5956
1·7256
0·721385
16
10·92
—
—
P
JF5-106f
142·4
309·3
0·4605
1·3339
0·719388
19
10·92
—
—
P
JF5-107
142·9
329·6
0·4336
1·2559
0·719048
16
14·27
—
JF5-108a
165·5
70·1
2·3624
6·8640
0·750770
18
29·87
JF5-108b
158·4
74·4
2·1295
6·1854
0·747588
27
JF5-108c
172·3
77·9
2·2121
6·4261
0·748871
11
JF5-108d
166·5
80·1
2·0784
6·0367
0·747189
JF5-108e
186·4
59·6
3·1293
9·1022
JF5-108f
178·1
64·3
2·7687
8·0492
JF5-108g
196·5
94·2
2·0852
JF5-109
208·7
95·3
2·1901
18·69q
18·26q
—
0·92
—
—
P
29·87
—
—
P
29·87
—
—
P
15
29·87
—
—
P
0·762043
20
29·87
—
—
P
0·756806
17
29·87
—
—
P
6·0565
0·747170
25
29·87
—
—
P
6·3623
0·749057
12
45·72
—
—
P
Localities 389 and 117, Sangerville Formation
JF389c-2
143·4
219·9
0·6519
1·8892
0·724577
16
P
JF389c-3
16·9
263·4
0·0641
0·1857
0·716500
21
C
JF389c-4
6·7
335·8
0·0199
0·0575
0·715819
27
C
JF389c-5
4·9
327·3
0·0151
0·0437
0·715882
20
C
JF389c-6
4·7
301·7
0·0156
0·0451
0·715932
20
C
JF389c-7
82·9
235·9
0·3516
1·0184
0·720290
11
C
JF389c-8
133·7
221·2
0·6044
1·7517
0·724689
14
P
JF389d
95·6
189·3
0·5050
1·4634
0·724026
17
P
JF117I
72·7
153·8
0·4724
1·3689
0·723923
15
P
1494
BICKLE et al.
Sample
Rb1
Sr1
Rb/Sr1
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
87
Rb/86Sr
87
Sr/86Sr
1r
Distance2 (m) d18O3
d13C
Rock type5
Waterville Formation Regional Pelite samples
JF411a
103·8
178·4
0·5818
1·6854
0·720325
12
P
JF674-5a
62·5
138·5
0·4511
1·3066
0·719054
23
P
JF674-5b
49·4
146·9
0·3361
0·9733
0·717429
12
P
159·2
113·0
1·4088
4·0867
0·734364
17
P
JF674a
Waterville Formation Limestone member
JF3A-1
42·6
901·8
0·0472
0·1366
0·709472
13
C
JF411-1
48·6
1109·9
0·0438
0·1268
0·712172
13
C
Hallowell Stock Contact
JF971-73
128·8
296·9
0·4338
1·2559
0·713848
13
G
JF971-78a 270·1
160·4
1·6839
4·8843
0·733341
16
G
JF971-78b 275·6
157·4
1·7510
5·0792
0·734196
16
G
JF971-78c
256·5
160·5
1·5981
4·6346
0·731423
21
G
JF971-78d 263·6
160·1
1·6465
4·7753
0·732449
43
G
JF971-78e 265·9
61·3
1·6485
4·7814
0·732895
17
G
Rb and Sr analyses accurate to ~±5% , Rb/Sr ratios to ±2%.
Distance measured with positive to west. Limestone–phyllite contact at 0·0 m on profile JF7 and at −0·20 m on profile JF5.
3
Oxygen and carbon isotope analyses on calcite except numbers with superscript q on quartz.
4
Samples collected at JF7 adjacent to contact about 25 m south of the main profile. These samples have not been included
on plots or fits below.
5
Rock types: P, pelite; C, carbonate; G, granite; numbers are fraction of pelite calculated from major element composition
as discussed in text.
1
2
XRF analysis at the Open University on an ARL 8420
spectrometer using a 3 kW Rh tube on fused discs made
with lithium metaborate–tetraborate flux and matrix
corrections calculated using the empirical Traill–
Lachance procedure.
TRACERS OF FLUID FLOW IN
METAMORPHIC ROCKS
Past fluid movement through porous media, such as
metamorphic rock, may be detected by advective displacements of chemical or isotopic compositions (e.g.
Bickle & McKenzie, 1987) or by the progress of mineralogical reactions driven by input of fluid out of equilibrium with the existing mineral assemblage (e.g.
Baumgartner & Ferry, 1991). Advective displacement of
chemical and isotopic tracers can only monitor flow
across compositional heterogeneities in rocks. As most
rocks are layered, this restricts study to the component
of flow perpendicular to layering, although some information on the relative fluxes along adjacent layers is
available from boundary-layer structure (e.g. Bickle &
Baker, 1990a; Bickle et al., 1995). Reaction progress
in fluid-composition-sensitive equilibria depends on the
amount and composition of fluid infiltrating the rock and
the amount of reaction is a measure of the infiltration
history. However, calculation of fluid fluxes from reaction
progress data requires a number of assumptions, the
validity of which may be questioned.
Most rocks exhibit too much chemical heterogeneity
and too little major element mobility for advective displacements of chemical tracers to be usefully resolved.
However, isotopic tracers often show large differences
between layers. 18O/16O ratios have been widely used
for monitoring fluid movement in metamorphic rocks
because oxygen partitions approximately equally into
fluid and solid phases. In addition, the partition is insensitive to fluid or rock composition. This allows detection of time-integrated fluid fluxes in excess of ~0·1
m3/m2 and whole-rock oxygen isotope compositions are
not sensitive to small amounts of subsequent fluid movement during retrogression. 13C/12C transport depends
on fluid composition (CO2 content), rock composition
(carbonate content) as well as the time-integrated fluid
flux. Rock 87Sr/86Sr ratios depend on the initial 87Sr/86Sr
ratio, the 87Rb/86Sr ratio and the amount and timing of
any changes in 87Sr/86Sr or 87Rb/86Sr ratio caused by
advection or diffusion in the fluid phase. Because Rb/
Sr ratios are controlled primarily by mineral modes
reflecting major element chemistry, the main effect of
advective or diffusive movement in the fluid phase is to
1495
JOURNAL OF PETROLOGY
VOLUME 38
NUMBER 11
NOVEMBER 1997
Table 2: List of symbols and constants
Variable
Definition
Units
A
Constant in permeability–porosity relationship [equation (11)]
C
Concentration, Cs solid, Cf fluid, C 1, C 2 initial concentrations across contact
Df
Diffusion coefficient (in fluid)
m2/s
Deff
Effective diffusivity of two-phase medium
m2/s
h
Scaling length used in transformation to dimensionless variables
m
Kd
Solid–fluid partition coefficient Cs /Cd
Kv
Fluid–solid partition coefficient by volume (qf/qsKd)
Pe
Peclet number
t
Time
x
Distance
s
z
Advective displacement of front
m
xo
Fluid (Darcy) or pore velocity
m/s
u
Porosity
q
Density, qs solid, qf fluid
Mg/m3
ri
One standard deviation error
s
Tortuosity coefficient for porosity
v2
Chi-squared parameter minimized in least-squares fits
Dh
C
18
13
Characteristic diffusion distance [equation (1)], DhSr, DOx
h , Dh Sr, d O and d C diffusion distances
modify 87Sr/86Sr ratios without significant changes to
87
Rb/86Sr ratios, and sampling on a variety of length scales
can give information on the timing of fluid movement as
well as on the amount of Sr-isotopic transport (e.g. Bickle
& Chapman, 1990). Below, we first discuss Rb–Sr isotopic
constraints on the timing of isotopic mobility, then the
oxygen, carbon and strontium isotope profiles.
ISOTOPIC TRANSPORT
Rb–Sr isochron systematics
Determination of Sr-isotopic mobility across lithological
boundary layers requires constraints on its timing to allow
correction for 87Rb decay. Rb–Sr isochron systematics are
completely reset over distances shorter than about 1/p
times the Sr-isotopic diffusion distance, DSrh taken as (see
Bickle et al., 1995)
DSrh=p
J
Dfust
qsKd/qf
(1)
where Df is the diffusion coefficient for Sr in the fluid, u
is porosity, s is tortuosity, qs and qf are solid and fluid
densities, and Kd is the solid–fluid partition coefficient
by mass for Sr (Table 2). The boundary-layer profile
modelled below indicates an Sr-isotopic diffusion distance
m
of ~2 m, which implies effective homogenization over
distances of <0·8 m.
Figure 2 illustrates isochron diagrams for ~1 cm slices
cut from a variety of whole-rock samples <0·15 m in
size from localities in the Waterville and Sangerville
Formations. For most of the samples the scatter about
the regression lines is little more than expected from
analytical error. Samples JF389c and -d comprise a set
of samples cut from a ~1 m profile across a 20 cm thick
calc-silicate band within phyllite at a diopside-grade
outcrop (Figs 1 and 2b). The Sr-isotopic systematics of
these samples confirms that Sr homogenization was nearly
complete on the metre scale, and homogenized carbonate
and phyllite which had significantly different Sr-isotopic
compositions when deposited (see Fig. 5 below). Figure 3,
a cumulative age diagram, shows that most of the thin
slice ages overlap within error of the mean age of granite
plutonism (Dallmeyer & Van Breeman, 1981, Fig. 4).
The deflection of regional metamorphic isograds around
the granite plutons is consistent with granite intrusion
being associated with the metamorphism (see Ferry,
1976). The small-scale Rb–Sr isochron systematics indicate that Sr-isotope mobility was also associated with
the metamorphism and that the Rb–Sr systematics have
remained largely undisturbed since then. As transport of
oxygen and carbon isotope variations must also depend
on advection and diffusion in a pore fluid, it seems
1496
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
Fig. 2. Isochron diagrams for ~1 cm slices cut from whole-rock samples from localities 5, 7 and 389. Errors quoted at 2r.
probable that mobility of these isotopic tracers took place
at the same time as for the strontium isotopes.
A second consequence of the ~2 m Sr-isotopic diffusion
distance is that samples averaged over length scales
greater than this should preserve their pre-metamorphic
Rb–Sr isotope systematics which will reflect the isotopic
composition of their source materials and any diagenetic
modifications. Figure 5 illustrates the Rb–Sr isotope systematics of the JF5 phyllite samples averaged over 5 m
intervals which regress about a line corresponding to an
age of 416±27 Ma consistent with their presumed Silurian depositional age. Figure 5 also shows that the
1497
JOURNAL OF PETROLOGY
VOLUME 38
Fig. 3. Cumulative diagram showing Rb–Sr isochron ages and 2r
error bars for centimetre-slice samples compared with the mean age
of granite plutonism from Rb–Sr whole-rock isochrons of Dallmeyer
& Van Breeman (1981).
Fig. 4. Isochron diagram of Hallowell Stock includes whole-rock data
of Dallmeyer & Van Breeman (1981) and six granite samples from
locality 971 adjacent to metasediment contact. Inclusion of these
samples does not significantly change either the age (previously 387±11
Ma) or the mean square weighted deviation (MSWD) of the isochron
fit.
Waterville limestone was not in Sr-isotopic equilibrium
with the phyllites at the time of deposition.
Boundary-layer isotopic profiles: locality 5
d18O profile: locality 5
Figure 6a illustrates the oxygen isotope profile across the
western margin of the ~100 m thick Waterville limestone
at locality 5 (Fig. 1). d18O values for calcites are plotted
NUMBER 11
NOVEMBER 1997
Fig. 5. Isochron diagram constructed from phyllite Rb–Sr isotopic
compositions averaged over 5 m intervals and compared with mean
composition of Waterville limestone samples collected away from layer
contacts.
for the limestone samples and one thin impure limestone
( JF5–104) in the phyllite. For the quartz separates, the
compositions that calcite would have had in equilibrium
with quartz have been plotted by subtracting 1‰ from
the measured quartz value to allow for calcite–quartz
fractionation as discussed below. In general, the profile
shows a transition from high d18O values in the limestone
to lower values in the phyllites consistent with the expected difference in original depositional values. At the
contact a number of both calcite and quartz isotopic
compositions scatter to either high or low values. These
samples are cut by a number of quartz and calcite veins.
Most of the veins are pre- or syn-metamorphic, and postmetamorphic veins are rare (see Rumble et al., 1991).
With the exception of the perturbed samples on the
contact, the profile shows a smooth transition from low
d18O in the phyllite to high d18O in the limestone with
the mid-point of the profile displaced ~2 m into the
limestone. This is characteristic of an initially sharp d18O
profile broadened by diffusion in the fluid phase and
displaced by advective transport of a component of the
fluid flux from phyllite to limestone (see Bickle & Baker,
1990a). The smooth transition in d18O from phyllite to
limestone suggests that fluid and solid were close to local
(grain-scale) equilibrium and that kinetic broadening of
the profile was limited. Significant kinetically limited
fluid–solid exchange would result in a step in d18O at
the contact. However, more restricted broadening of
fronts by kinetically limited fluid–solid exchange is impossible to distinguish from diffusive broadening (Baker
& Spiegelman, 1995) unless upper and lower boundary
layers may be compared (Skelton et al., 1997a). The
coupling of oxygen, carbon and strontium advection and
diffusion distances discussed below, assuming all the front
1498
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
Fig. 6. (a) d18O profile across the western margin of the Waterville limestone at locality 5. Open symbols denote samples not included in leastsquares fits. Continuous curve is least-squares fit to advective–diffusive transport model for uniform flow from phyllite to limestone with
composition profile as modelled in (b) by equation (5) and the parameters given for this fit with errors at 1r. z∗ is advective displacement and
DhOx diffusion distance as defined by equation (1). Dashed line is least-squares fit for pinned boundary condition (see text). (b) Fourier series fit to
carbonate/(carbonate + pelite) calculated from major element composition of samples. [Note that profile is padded with data (not shown) at
<−5 m and > + 5 m to damp smaller wavelength terms where data are sparse away from contact.]
broadening results from diffusion, produces comparable
fluid XCO2 and Sr concentrations and suggests that a
significant part of the broadening of each front (~50%)
is related to diffusion, as kinetically limited fluid–solid
exchange is likely to produce substantial decoupling of
transport distances (see Bickle, 1992). Below, we describe
the geometry of the fronts by ascribing all the broadening
to diffusive processes. The most important implication
of the isotopic boundary-layer studies for the subsequent
discussion of the relationship between reaction progress
and fluid flow is the distance over which isotopic and
chemical species have been homogenized. The precise
mechanism of this homogenization is of second-order
importance and a description of the boundary layers by
advection and diffusion provides a good measure of
homogenization distances for the various isotopic species
studied.
One-dimensional advective and diffusive transport of
a tracer such as d18O is described by a differential
equation of the form (e.g. Bickle & McKenzie, 1987)
C
D
∂C
∂2C
∂C qsK d
(1−u)+u +xou =Deff 2
∂t qf
∂x
∂x
(2)
where C is the concentration of a chemical tracer or an
isotope ratio, Deff is the effective bulk diffusion coefficient
(Df u s), xo is the fluid (Darcy) velocity and t is time
(Table 2). For tracers such as d18O in metamorphic rocks
Kd, the solid–fluid partition coefficient for the element, is
large compared with porosity, u, and the first term in
equation (2) can be simplified by the approximation
qK
qsKd
(1−u)+u≈ s d .
qf
qf
(3)
The profile in Fig. 6 has been modelled to calculate
advective and diffusive transport distances by solving
equation (1) for appropriate boundary conditions. The
calculation requires (1) an initial shape of the isotopic
profile across the contact, (2) the value of calcite–quartz
oxygen isotope fractionation and (3) criteria for excluding
sample points whose compositions have been perturbed
by events younger than the syn-metamorphic fluid flow.
The initial (pre-metamorphic) shape of a d18O profile
may be determined by comparison with similar rocks
which have not undergone significant isotopic transport
(e.g. Bickle & Baker, 1990a). In this study very low grade
equivalent strata are not available and data from the
chlorite-grade locality 7 are discussed below. We have
estimated a pre-metamorphic d18O profile by analysing
the major element compositions of the rocks and modelling the lithological variation in terms of mixtures
between end-member limestone and phyllite rocks. The
major element compositions are listed in Table 3. An
average of samples JF5-89 and -90 was taken as the type
limestone and JF5-105, -106 and -107 as the type phyllite.
The major elements SiO2, Al2O3, total iron, CaO and
K2O differ significantly between these end-members and
a least-squares technique was used to calculate the bestfit end-member proportions in each sample by minimizing the sum of the deviations squared weighted by
1499
JOURNAL OF PETROLOGY
VOLUME 38
NUMBER 11
NOVEMBER 1997
Table 3: Major element analyses
Sample:
JF5-79b
JF5-80b
JF5-81a
JF5-82a
JF5-83
JF5-84b
JF5-86
JF5-87
JF5-88
JF5-89
SiO2
TiO2
Al2O3
Fe2O3∗
FeO
MnO
MgO
CaO
Na2O
K 2O
P 2 O5
LOI
Total
53·17
0·860
15·9
0·56
5·9
0·27
3·88
9·1
0·69
3·29
0·243
4·65
99·18
53·71
0·800
15·7
0·79
5·62
0·30
4·05
8·97
0·8
3·16
0·218
4·76
99·51
56·85
0·891
18·47
0·79
6·48
0·19
3·55
5·21
1·52
3·3
0·206
1·34
99·51
61·73
0·784
16·85
0·83
5·69
0·13
2·66
3·63
1·45
3·29
0·129
1·69
99·49
48·33
0·911
19·18
7·74†
0·25
3·47
10·89
1·10
2·79
0·14
3·78
98·58
25·49
0·388
8·32
0·29
3·54
0·1
3·51
30·31
0·94
1·71
0·057
24·36
99·5
26·91
0·367
7·78
0·24
3·18
0·11
3·57
29·8
1·07
1·64
0·056
24·07
99·15
35·54
0·531
11·37
0·38
4·72
0·14
4·21
21·46
1·07
2·52
0·084
16·4
98·96
26·65
0·296
5·45
0·18
2·29
0·09
3·38
31·51
1·19
0·96
0·055
27·07
99·37
27·7
0·266
4·75
0·13
2·02
0·11
3·13
32·3
1·15
0·78
0·05
27·16
99·76
Sample:
JF5-90
JF5-91a
JF5-91b
JF5-91c
JF5-91d
JF5-91e
JF5-91f
JF5-91g
JF5-92C
JF5-93
SiO2
TiO2
Al2O3
Fe2O3∗
FeO
MnO
MgO
CaO
Na2O
K 2O
P 2 O5
LOI
Total
28·43
0·255
4·31
0·16
1·64
0·06
3·13
31·89
1·24
0·56
0·049
27·82
99·73
19·88
0·238
4·36
0·26
1·58
0·07
3·32
37·1
1·04
0·72
0·048
31·5
100·3
23·44
0·219
3·66
0·1
1·48
0·06
3·56
34·7
1·12
0·48
0·046
30·87
99·9
24·9
0·216
3·84
0
1·52
0·06
3·23
34·24
1·18
0·52
0·045
29·82
99·74
26·77
0·588
13·48
0·43
4·42
0·07
7·46
22·09
1·49
3·51
0·078
17·19
98·07
23·43
0·364
8·4
0·25
2·57
0·08
4·36
31·16
1·48
1·75
0·047
24·74
98·92
21·4
0·186
3·14
0·24
1·31
0·08
2·28
37·86
0·72
0·52
0·046
31·36
99·29
26·83
0·328
5·94
0·22
2·22
0·06
4·16
30·92
1·02
1·42
0·051
25·68
99·1
26·71
0·347
7·32
0·3
2·85
0·11
3·83
30·36
1·05
1·59
0·045
24·27
99·1
18·36
0·199
3·47
0·25
1·84
0·27
2·46
38·24
1·02
0·51
0·039
32·44
99·3
Sample:
JF5-94
JF5-95
JF5-96
JF5-99
JF5-101
JF5-102
JF5-104c
JF5-105
JF5-106a
JF5-107
SiO2
TiO2
Al2O3
Fe2O3∗
FeO
MnO
MgO
CaO
Na2O
K 2O
P 2 O5
LOI
Total
22·54
0·305
6·56
0·27
2·46
0·07
3·87
33·01
0·84
1·31
0·044
27·51
99·06
20·1
0·271
5·6
0·3
2·68
0·21
3·59
34·98
0·77
1·12
0·053
29·1
99·1
46·65
1·285
25·5
1·19
6·36
0·19
3·16
2·47
1·25
7·01
0·362
3·2
99·33
65·48
0·871
14·72
0·6
5·13
0·21
2·73
3·58
1·6
2·55
0·192
1·21
99·45
63·08
1·008
16·43
0·68
6·6
0·13
2·35
1·94
1·18
4·03
0·148
2·06
99·64
56·45
0·788
14·23
0·73
5·69
0·25
3·43
7·95
1·03
3·13
0·254
4·96
99·53
65·4
0·944
17·21
1·08
5·01
0·11
1·76
0·32
0·62
4·31
0·11
2·59
100·02
53·38
0·795
17·86
2·4
9·06
2·54
4·03
3·88
1·5
3·26
0·121
0·95
100·79
54·37
0·775
17·35
2·45
9·03
2·46
3·94
3·35
1·43
3·29
0·143
1·17
100·77
21·16
0·318
6·8
0·26
2·63
0·07
4·00
32·82
0·54
1·72
0·05
28·56
99·21
Av. pelite‡
SiO2
Al2O3
Fe2O3
CaO
K 2O
Sr and 87Sr/86Sr
57·7
17·47
10·53
2·52
3·62
252
Av. limestone‡
28·1
4·53
2·18
32·1
0·67
0·71290 734
Percentage error
0·3
0·3
0·5
0·5
1·0
0·70881
∗Fe2O3 determined by difference between total Fe2O3 by XRF and FeO by titration.
†Total iron as Fe2O3.
‡Adopted in least-squares fits to determine proportion of pelite and limestone end-members.
1500
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
a factor inversely proportional to analytical error. The
results are illustrated in Fig. 6b.
Advective–diffusive broadening of the d18O profile by
uniform flow across the contact at locality 5 has been
calculated by fitting a Fourier series to the data in
Fig. 6b and using a least-squares minimization routine
to determine the advective and diffusive distances and
the initial compositions of limestone and phyllite away
from the contact (Fig. 6a). The data in Fig. 6b were
padded with additional points (C=0) at >5 m and (C=
1) <–5 m to eliminate aliasing into high-frequency terms
away from the contacts where sample spacings are larger.
The initial profile, determined by an overdetermined
linear least-squares fit, is thus expressed as
k
Cx=A1+; [A2i cos(pix′ )+A2i+1sin(pix′ )]
(4)
i=1
where the terms Ai are the Fourier coefficients and x′ is
dimensionless distance (distance x=hx′ where h is the
scaling length). The curve in Fig. 6b was calculated with
27 Fourier terms and with dimensionless distance scaled
to h=60 m. The equation of the advectively–diffusively
modified profile is given by [see Bickle & McKenzie
(1987), equations (1) and (11); Carslaw & Jaeger (1959),
p. 93]
k
Cx=A1+; [(A2i cos(pi[x′−z∗′])
i=1
+A2i+1sin(pi[x′−z∗′]))e−pi t ′ ]
2 2
(5)
where t′ (dimensionless time) and z∗′ (dimensionless
advective distance) are given by
qK
h2 s d
qf
t′.
t=
Deff
(6)
and
xout
qsKd
qf
z∗=z∗1h=
(7)
Equation (2) has been solved for both the solution above
[equation (5)] and for a pinned boundary condition with
an initial step function (Bickle & Baker, 1990a). The
parameters, C 1, the initial d18O of pure limestone, C 2,
the initial d18O of pure phyllite, z∗′, the advective displacement and t′, dimensionless time were estimated by
minimizing v2 over the n data points by the non-linear
Levinson–Marquhardt method (Press et al., 1986), where
n
v2=;
i=1
C
(Cm−Cc)2
r2i
D
(8)
and where Cm is the measured d18O, Cc is the calculated
d18O and ri is the estimated error on the analysed isotope
composition. The initial d18O (Fig. 6b) is assumed directly
proportional to the carbonate and phyllite proportions
calculated from the major element chemistry with the fit
forced to pure limestone or pure phyllite away from the
contact region.
Peak metamorphic temperatures at locality 5 were
~460°C [Ferry (1994) and as discussed below]. Quartz–
calcite oxygen isotope fractionation (Dqtz-cc) at 450°C was
estimated from experiment to be 0·75 by Matthews (1994)
and 0·8 by Clayton & Kieffer (1991), but 1·6‰ from an
empirical calibration by Sharp & Kirschner (1993). The
minimum analytical uncertainty is ±0·25‰. Four
samples from which both calcite and quartz were analysed
give fractionation factors of 0·14, 1·01, 1·49 and 1·58
with a mean of 1·0. Two of these samples ( JF5–80b
and JF5–83, with fractionations of 1·01 and 0·14) were
collected close to the contact, where late veining has
perturbed d18O values. Intrinsically, it should be possible
to solve for quartz–calcite d18O fractionation, in addition
to the other four unknown parameters, by least-squares
minimization. However, the minimization becomes less
robust as the number of parameters is increased. Models
were run with Dqtz–cc between 0·5 and 2·0. The models
with Dqtz-cc between 1·0 and 1·5 gave the best fits with
rox, the standard deviation of data points about the bestfit curve, ~0·2‰. The standard deviation about the bestfit curve increases markedly for values of Dqtz-cc <1·0 or
>1·5 with rox increasing to ~0·33 (Dqtz-cc=0·5 or 2·0).
Models are evaluated for Dqtz–cc=1·0. The difference
between this value and the best estimate from the experimental determinations is probably not significant
given the precision of the experiments and the possibility
of post-fluid flow exchange. The metamorphic temperature is close to the blocking temperature for quartz
(Eiler et al., 1992) but calcite may have exchanged oxygen
isotopes with the coexisting calc-silicate minerals during
cooling.
Some samples adjacent to the contact contain quartz
and calcite veins. These may relate to post-peak metamorphic fluid flow and four quartz separates have d18O
values significantly less, and one more, than the other
quartz or calcite values. Thus samples JF5-83, -97, -98,
-100 and -102 have been excluded from the fitting
routine. The best fit of the 24 remaining data points to
equation (5) is shown in Fig. 6. The 24 points scatter
about the best-fit line with a standard deviation (rox) of
0·24‰, which is close to analytical precision. The advective displacement is –1·99±0·41 m (1r errors), which
implies a time-integrated fluid flux of 3·2±0·7 m [equation (7), Kd for oxygen is ~1·6]. The diffusion distance,
DOx
h [equation (1)], is 6·35±0·26 m. Figure 6b shows that
lithological broadening of the initial profile is not more
than ~2 m. Assumption of an initial step-function profile
1501
JOURNAL OF PETROLOGY
VOLUME 38
gives a similar quality fit (rox=0·21‰), with a timeintegrated fluid flux of –3·5±0·8 m and a diffusion
distance (7·6±2·0 m) within error of the values modelled
to fit the initial profile calculated from the lithological
data. This insensitivity to the relatively small initial broadening of the profile is to be expected because the total
diffusion distance varies as the square-root of the sum of
the squares of individual events (see Bickle et al., 1995).
To increase a profile diffusively broadened by 2 m to a
profile equivalent to that with a diffusion distance of
6·2 m requires an event with a diffusion distance of 5·8 m.
A least-squares fit with a pinned boundary condition
to simulate a large layer-parallel flux in the phyllite such
that the oxygen isotope composition at the contact is
buffered to a constant value (see Bickle & Baker, 1990a),
fits only slightly less well (dashed curve in Fig. 6a), with
a mean deviation about the line of 0·27‰, and gives
similar values for advective displacement (z∗=
–2·0±2·7 m, DOx
h =10·6±6·2 m, 1r errors). The similarity between the uniform flow and pinned boundary
condition fits is to be expected at an upstream boundary
where advection transports the boundary layer structure
a distance comparable with the diffusion distance (see
Bickle & Baker, 1990a). The structure of the isotope
profile preserves little information about the relative
flow conditions in the phyllite vs limestone and this
information should be sought at the downstream contact
of the Waterville limestone member which is not exposed
at locality 5.
d13C profile at locality 5
The d13C profile is illustrated in Fig. 7. This shows
a decrease in d13C within the limestone towards the
limestone–phyllite contact similar to the d18O profile but
over a much shorter distance. Modelling the d13C profile is
more problematic than the oxygen-isotope profile because
only one carbonate value is available from the phyllite
away from the limestone contact (sample JF5–104, a thin
impure limestone) and calculation of the initial profile
from the lithological data requires an assumption about
the initial carbonate content of the phyllite end-member.
However, the most impure carbonate sample closest to
the contact ( JF5–83 at 3 cm into the limestone) contains
28% of the carbonate component and if the phyllite
component contained 5% carbonate (average pelitic sediment) then this sample would have >80% of its carbonate
from the limestone end-member. For these reasons the
most appropriate initial profile for modelling the d13C
transport is a step function.
Figure 7 illustrates the least-squares fit for a uniform
flow solution to equation (2) (Bickle & Baker, 1990a) to
15 of the 16 samples with calcite d13C analyses. The data
scatter about the best-fit line with a standard deviation
NUMBER 11
NOVEMBER 1997
Fig. 7. Fit to calcite d13C compositions at locality 5 assuming uniform
flow across contact. Μ, point excluded from fit.
of 0·36‰, and give an advective displacement of
–0·42±0·09 m and a diffusion distance of 1·64±0·32 m.
Given the uncertainty over the appropriate boundary
conditions, the significance of the fit parameters is unclear.
87
Sr/86Sr profile at locality 5
The post-Acadian 380 Ma 87Sr/86Sr profile is shown in
Fig. 8a. This is calculated from the present-day 87Sr/86Sr
and 87Rb/86Sr ratios and is inferred to represent the Srisotopic composition immediately after fluid-flow events
associated with the Acadian metamorphism because the
small-scale Rb–Sr isochrons give ages within error of
380 Ma (Fig. 2). Also shown in Fig. 8a is an estimate of
the 87Sr/86Sr profile immediately before the Acadian
metamorphism at 380 Ma. For the phyllite this is calculated assuming that the samples had the age and initial
87
Sr/86Sr ratio given by the >5 m scale isochron (Fig. 5)
and that their Rb/Sr ratios remained unchanged during
the metamorphism. For the limestone samples the initial
ratio is calculated assuming that the least radiogenic
samples in the centre of the horizon (initial 87Sr/86Sr at
416 Ma of 0·7085) are representative of the pre-Acadian
Sr-isotope systematics. The small Rb/Sr ratios of the
limestone samples make estimates of pre-Acadian 87Sr/
86
Sr ratios insensitive to both age corrections and plausible
changes in 87Rb/86Sr ratios. The Sr-isotopic compositions
have been calculated presuming that they comprise a
mixture of limestone and phyllite components calculated
from the whole-rock compositions as in Fig. 6b and that
the limestone end-member contains 2·9 times the amount
of Sr in the phyllite end-member. This profile (Fig. 8b)
indicates that the initial profile is broadened from a step
1502
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
Fig. 8. (a) Pre-Acadian 380 Ma 87Sr/86Sr profile (Β) calculated from assumed initial 87Sr/86Sr and post-Acadian 380 Ma 87Sr/86Sr profile (Χ)
calculated from present-day 87Sr/86Sr and Rb/Sr ratios. Line is least-squares fit assuming uniform flow and step-function initial profile. (b) Detail
of Pre-Acadian 380 Ma profile at contact showing broadening about step-function is equivalent to a diffusion distance of less than ~0·5 m.
Errors given at 1r.
function by the equivalent of less than ~0·5 m diffusion
distance.
Figure 8a also shows a least-squares fit to the postAcadian 87Sr/86Sr profile assuming uniform flow across
the contact and an initial step-function profile [see Bickle
& Baker (1990a), equation (10)]. This gives an advective
displacement of –0·62±0·22 m and a diffusion distance
of 2·17±0·77 m. If the profile had an initial broadening
equivalent to a diffusion distance of 0·5 m, this would
only decrease the estimated diffusion distance by 0·06 m.
Isotopic constraints on fluid composition
and movement at locality 5
The three isotopic profiles are all consistent and indicate
advective displacements from the phyllite to the carbonate. The oxygen-isotope volume partition coefficient
(Kv) is ~0·6 and advective displacement of the oxygenisotope front implies a time-integrated fluid flux of
3·2±0·7 m3/m2 (1r error) perpendicular to the phyllite–
limestone contact. Transport of the carbon and strontium
isotope fronts depends on fluid composition (CO2 or Sr
content) in addition to the time-integrated fluid flux.
Given the time-integrated fluid flux and rock compositions, it is possible to calculate the fluid CO2 and Sr
contents from their advective displacements by solving
equation (7). Similarly, given the diffusion distance for
oxygen, independent solutions for fluid CO2 and Sr
contents may be obtained from equation (1) assuming
DSrf=0·5DOxf (Bickle & Chapman, 1990). Calculated fluid
CO2 and Sr concentrations are listed in Table 4.
Table 4: Fluid compositions
XCO
2
1r
Sr (p.p.m.)
1r
Advection
0·06
0·02
400
164
Diffusion
0·02
0·007
75
27
Both the estimates of fluid XCO2 and fluid Sr concentration are higher from advective transport than from
diffusive transport although both overlap at 2r, given
the rather large errors. The XCO2 fluid compositions are
relatively water rich, consistent with the estimate by Ferry
(1987, 1994) that XCO2 ~0·10 at 430°C, given that the
significant fluid flow occurred while the rocks were heated
from ~400°C to 460°C. The mineral equilibria discussed
below imply peak metamorphic temperatures of
465±5°C and XCO2 between 0·25 and 0·5. Sr concentrations in metamorphic fluids are poorly constrained.
As Sr is probably complexed by chloride (e.g. Palmer,
1992), the relatively high Sr concentrations imply relatively saline fluids by comparison with fluids from black
smokers at mid-ocean ridges (Sr ~24 p.p.m. in 1 molar
chloride fluid) and fluids from the Salton Sea geothermal
system (Sr ~440 p.p.m. in ~6 molar chloride fluid;
Helgeson, 1967).
Isotopic boundary layers at locality 7
Rb/Sr, oxygen, carbon and strontium isotope profiles
across the eastern contact of the Waterville limestone at
1503
JOURNAL OF PETROLOGY
VOLUME 38
locality 7 (Fig. 1) are shown in Figs 9–12. The major
element composition of samples across the contact is not
available at this locality but a plot of Rb/Sr ratios (Fig. 9)
indicates that the transition from carbonate to phyllite
takes place over <0·5 m. The d18O profile (Fig. 10) shows
the expected drop in values as the phyllite contact is
approached. Silicate analyses are not available, but if the
calcite-equivalent value for unaltered phyllite of 17·46 is
assumed (Fig. 6), a least-squares fit for uniform flow
with an initial step-function profile implies an advective
displacement of +0·9±0·5 m and a diffusion distance,
DOx
h , of 1·4±4·0 m (2r errors). Despite the relatively large
errors, both the advective distance and diffusion distance
are significantly less than at locality 5 (2·0 and 6·4 m,
respectively). The d13C profile (Fig. 11) is unconstrained
in the phyllite. A fit assuming d13C=−6·75‰ (as at
locality 5) gives an advective distance of −1·0±0·7 m
and a diffusive distance of 2·8±2 m (but note the large 2r
errors). The 380 Ma 87Sr/86Sr profile implies a diffusion
distance of 0·7±0·5 m and no detectable advective displacement (0·05±0·13 m) (Fig. 12). The diffusion distance is again significantly smaller than that at locality 5
(2·2 m). The smaller diffusion and advective displacements are consistent with lower fluid fluxes and
lower temperatures at locality 7 compared with locality
5. As the temperature difference between localities 5 and
7 is only ~60°C (e.g. Ferry, 1986) the main control on
the different diffusion distances is probably the porosity
and the time the porosity was open. Diffusion distance
scales as the square-root of the product of porosity
and time [equation (1)] and the factor of three to four
difference in diffusion distances implies that the timeintegrated fluid flux was an order of magnitude greater
in the higher-grade Waterville limestone at locality 5.
This is consistent with the change in reaction progress
between the two localities in which most samples of the
Waterville limestone at locality 7 have undergone little
or no carbonate breakdown but samples from locality 5
contain a significant fraction of calc-silicate minerals
(Ferry, 1987).
REACTION PROGRESS AND
ISOTOPIC BOUNDARY LAYERS
Ferry (1987, 1988, 1994) and Baumgartner & Ferry
(1991) observed large differences in reaction progress
from layer to layer on a scale of <0·5 m in the Waterville
limestone, from which they calculated order-of-magnitude variations in the time-integrated fluid flux over
the same length scale. The time-integrated fluid flux was
calculated from progress of decarbonation reactions in
the limestone driven by infiltration of more water-rich
fluids. Their model for calculating time-integrated fluid
NUMBER 11
NOVEMBER 1997
Fig. 9. Rb/Sr ratios across eastern margin of Waterville limestone at
chlorite-grade locality 7. Transition from carbonate to phyllite takes
place over <0·5 m.
Fig. 10. d18O profile across eastern margin of Waterville limestone at
chlorite-grade locality 7 (Fig. 1). Least-squares fit made by fixing d18O
value of calcite in phyllite to the same value as estimated at locality 5
(17·46‰). Errors at 1r.
fluxes is based on the observation that isobaric nearunivariant calc-silicate assemblages with substantial reaction progress are stable across significant distances on
the ground. Given near-equilibrium fluid–solid exchange,
such an array of internally buffered calc-silicate assemblages is only possible if fluid flow is up temperature
and largely layer parallel. Baumgartner & Ferry (1991)
and Ferry (1994) calculated the time-integrated fluid flux
assuming flow is horizontal, up a fixed temperature
gradient given by the regional metamorphic gradient
which then specifies a gradient (∂XCO2/∂z) in XCO2 for a
given univariant assemblage. The time-integrated fluid
flux qm (in mol/m2) is given by
nCO2−XCO2 (nCO2+nH2O)
∂XCO2/∂z
qm=
(9)
where nCO2 and nH2O are the CO2 and H2O (mol/m3)
1504
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
homogenized over a distance of ~6 m for d18O, ~1·6 m
for d13C and ~2 m for 87Sr/86Sr. Carbon isotopes are
transported as CO2 species in the fluid and therefore
the homogenization distances for CO2 species and d13C
should be identical. If the isotope homogenization took
place during the same period as the fluid infiltration,
gradients in XCO2 should be smoothed over a distance of
~1·6 m and reaction progress in centimetre-scale layers
with assemblages stable at higher XCO2 for an equivalent
amount of reaction progress should be driven to higher
amounts of reaction progress than in adjacent layers with
lower reaction progress at the same XCO2.
Fig. 11. d13C profile at locality 7. Least-squares fit for uniform flow
and initial step function has been forced through d13C=−6·75‰ in
phyllite. Errors at 1r. Open symbols indicate samples not included in
fit.
Fig. 12. 380 Ma 87Sr/86Sr profile across margin of Waterville limestone
at locality 7. Least-squares fit for uniform flow and initial step function.
Errors at 1r. Open symbols indicate samples not included in fit.
produced by the reaction. In the Waterville limestone
at locality 5, Baumgartner & Ferry calculated timeintegrated fluid fluxes between ~200 and 5000 m3/m2,
and samples used in this study imply a similar range of
values as discussed below. The isotopic boundary-layer
profiles record cross-layer flow of ~3 m3/m2, which is a
factor of 102−103 less than the layer-parallel flow inferred
from reaction progress. The cross-layer time-integrated
fluid flux is also insufficient to cause the observed decarbonation of the ~100 m thick Waterville limestone,
which would require ~50 m3/m2 of a pure H2O-rich fluid
calculating the time-integrated fluid flux from equation (5)
of Bickle & Baker (1990b), assuming that the biotiteforming reaction produces 2·9 mol CO2/l and buffers
fluid composition to XCO2=0·1. This is a minimum
estimate, as the infiltrating fluid is likely to have contained
significant CO2. However, the isotopic profiles were
Estimates of pressure, temperature and
fluid composition
We have investigated the relationship between reaction
progress and XCO2 for samples used in this study.
Waterville limestone samples contain the maximum
phase assemblage quartz–muscovite–biotite–chlorite–
plagioclase–calcite–ankerite–rutile–iron sulphide. Subareas of some samples lacked either chlorite or muscovite.
Phyllitic rocks contain combinations of quartz, muscovite,
biotite, chlorite, calcite, plagioclase, garnet, ilmenite, rutile and sulphides. Calcite is absent from the matrix of
most of the phyllitic samples (although often present in
late veins) and garnet is present only in some of the
samples. Microprobe analyses are listed in Table 5. The
analysed phases are relatively homogeneous, with the
exception of plagioclase, which has a limited range of
anorthite contents.
The phases quartz, muscovite, biotite, chlorite, plagioclase, calcite and ankerite in the locality 5 samples
contain 14 end-members for which thermodynamic data
are available and have activities in the phases which give
consistent pressure and temperature estimates. Average
pressure and temperature are calculated using
THERMOCALC (Holland & Powell, 1990). Activities
of components in muscovite and phlogopite are calculated
as given by Holland & Powell (1990) using activity
coefficients of Eugster et al. (1972), dolomite is calculated
assuming two-sites ideal mixing and calcite from one-site
ideal mixing, plagioclase after model 1 of Holland &
Powell (1992) and chlorite after Holland & Powell
(1990). All the independent reactions from all the rocks
which contain the assemblage quartz–muscovite–
biotite–chlorite–plagioclase–calcite–ankerite give XCO2dependent results with P=8·9±4·0 kbar and T=
558±35°C (2r error estimates) at XCO2=0·1 to P=
2·9±1·2 kbar and T=455±40 at XCO2=0·5. Ferry
(1980, 1986) estimated pressure to have been 3·5±0·2
kbar and the temperature 450°C at locality 5 from a
variety of biotite–garnet–plagioclase–aluminium silicate
and calcite–dolomite equilibria. The Waterville limestone
1505
JOURNAL OF PETROLOGY
VOLUME 38
NUMBER 11
NOVEMBER 1997
Table 5: Electron microprobe mineral analyses recalculated as cations per molar unit
Muscovite
K
Na
Ca
Fe
Mg
Mn
Ti
Al
Si
Oxide sum
K/(K+Na)
Biotite
K
Na
Ca
Fe
Mg
Mn
Ti
Al(VI)
Al(IV)
Si
Oxide sum
Fe/(Fe+Mg)
Calcite
C
Mg
Fe
Mn
Oxide sum
Ankerite
Ca
Mg
Fe
Mn
Oxide sum
Fe/(Fe+Mg)
Plagioclase
Ca
Na
K
Al
Si
Oxide sum
Max. X an
Range X an
Chlorite
Fe
Mg
Mn
Ti
Al
Si
Oxide sum
Fe/(Fe+Mg)
JF5-79a
JF5-99
JF5-91a
JF5-91b
JF5-91c
JF5-91d
JF5-91e
0·8821
0·0674
0·0024
0·0818
0·0674
0·0000
0·0241
2·7170
3·0948
94·79
0·9290
0·8462
0·0811
0·0044
0·0668
0·0657
0·0017
0·0195
2·7833
3·0924
96·00
0·9125
0·9004
0·0567
0·0028
0·0443
0·1532
0·0000
0·0209
2·6612
3·1430
95·15
0·9408
0·8865
0·0584
0·0021
0·0526
0·1708
0·0021
0·0230
2·6596
3·1333
95·85
0·9382
0·8943
0·0575
0·0063
0·0473
0·1733
0·0003
0·0220
2·6367
3·1485
95·14
0·9396
0·8825
0·0644
0·0080
0·0407
0·1678
0·0000
0·0215
2·6265
3·1639
94·83
0·9320
0·8748
0·0608
0·0055
0·0443
0·1578
0·0007
0·0186
2·6757
3·1364
95·44
0·9350
0·8836
0·0605
0·0025
0·0473
0·1518
0·0011
0·0165
2·6709
3·1422
95·74
0·9359
0·8544
0·0171
0·0113
0·9774
1·2854
0·0096
0·0910
0·4729
1·2189
2·7811
96·07
0·4320
0·8862
0·0199
0·0052
1·2695
1·0409
0·0179
0·0832
0·4526
1·2624
2·7376
95·95
0·5494
0·8718
0·0105
0·0080
0·5632
1·8255
0·0014
0·0613
0·4028
1·1309
2·8691
96·15
0·2358
0·8767
0·0096
0·0113
0·5863
1·7157
0·0014
0·0699
0·4443
1·1293
2·8707
95·21
0·2548
0·8624
0·0094
0·0171
0·5588
1·8318
0·0008
0·0635
0·3940
1·1235
2·8765
95·87
0·2337
0·8682
0·0160
0·0039
0·4903
1·9115
0·0008
0·0635
0·3959
1·1389
2·8611
96·09
0·2041
0·8250
0·0144
0·0058
0·5046
1·8748
0·0014
0·0571
0·4210
1·1053
2·8947
95·61
0·2120
0·8789
0·0165
0·0099
0·5217
1·8461
0·0008
0·0613
0·4146
1·1439
2·8562
95·86
0·2203
0·9053
0·0395
0·0295
0·0275
56·43
0·8410
0·0410
0·0452
0·0642
56·31
0·9574
0·0306
0·0113
0·0006
56·18
0·9516
0·0354
0·0119
0·0010
56·03
0·9500
0·0368
0·0121
0·0012
56·68
0·9595
0·0293
0·0092
0·0020
55·64
0·9523
0·0350
0·0111
0·0016
56·27
0·9601
0·0288
0·0097
0·0014
56·45
no ank
no ank
1·0153
0·8633
0·1190
0·0024
53·05
0·1211
1·0168
0·8688
0·1120
0·0025
52·65
0·1142
1·0128
0·8628
0·1223
0·0020
52·99
0·1241
1·0218
0·8592
0·1129
0·0061
56·02
0·1162
1·0169
0·8646
0·1123
0·0061
52·75
0·1150
1·0192
0·8630
0·1137
0·0041
52·68
0·1164
0·8760
0·1160
0·0000
1·8880
2·1130
100·06
0·8840
0·83–0·88
0·4390
0·5680
0·0000
1·4430
2·5500
100·41
0·4410
0·41–0·44
0·3050
0·6920
0·0000
1·3060
2·6940
99·65
0·3060
0·29–0·31
0·2710
0·7210
0·0000
1·2730
2·7290
99·98
0·2740
0·24–0·27
0·2714
0·7250
0·0000
1·2616
2·7366
99·77
0·2679
0·2700
0·5910
0·4130
0·0000
1·5870
2·4110
99·98
0·5885
0·31–0·59
0·3840
0·6220
0·0000
1·3900
2·6070
99·52
0·3863
0·27–0·39
1·7776
2·6718
0·0280
0·0074
2·7265
2·7038
88·46
0·3995
2·2444
2·1551
0·0481
0·0123
2·7956
2·6613
88·44
0·5102
1·0703
3·4811
0·0018
0·0035
2·6292
2·7363
86·82
0·2353
1·0154
3·4965
0·0053
0·0091
2·6240
2·7626
87·05
0·2251
0·9930
3·5616
0·0000
0·0025
2·6212
2·7433
87·23
0·2180
0·8323
3·7513
0·0007
0·0014
2·5904
2·7612
86·75
0·1816
0·8775
3·6691
0·0021
0·0046
2·6250
2·7514
87·39
0·1930
1506
JF5-91f
0·3500
0·6470
0·0000
1·3430
2·6540
100·75
0·3480
0·28–0·35
0·9226
3·6138
0·0032
0·0032
2·6418
2·7398
87·14
0·2034
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
Fig. 13. Fit parameter, f, from Holland & Powell (1990) as a function
of XCO2 for P=3·5 kbar and for average temperature calculation for
15 end-members in quartz–muscovite–biotite–chlorite–plagioclase–
calcite–ankerite assemblage from JF5–91f and JF5–79a (less ankerite).
f values of <1·45 imply fit within 95% confidence level given errors
on thermodynamic variables and activities of end-members.
assemblages do not put good constraints on either pressure or XCO2, as can be seen in Fig. 13, which shows the
goodness of fit parameter, f (rfit), as a function of XCO2 at
3·5 kbar for sample JF5–91a. The eight independent
equilibria yield temperature estimates within error expected from the uncertainties on the thermodynamic
data set and from the end-member activities (propagated
from analytical uncertainties) if f<1·45 (see Holland &
Powell, 1990). Figure 13 shows that this condition is
satisfied for 0·25 < XCO2< 0·5 for all 15 end-members
and for a pressure of 3·5 kbar at temperature between
460 and 470°C. Exclusion of the least well fitting endmember, celadonite, makes no difference to the range of
acceptable fits or their temperatures. Exclusion of dolomite extends the range of acceptable fits to fluid compositions with XCO2 as low as 0·02. All the limestone
samples record similar temperature and XCO2 estimates
at 3·5 kbar. The mean temperature estimate (465±10°C)
is very close to the estimate by Ferry (1986) of 450°C,
and identical to estimates by Ferry (1994). Phyllite samples
which lack ankerite give a slightly larger range of fluid
XCO2 contents at similar temperatures and acceptable f
values (e.g. JF5–79a, Fig. 13).
Reaction progress
Reaction progress has been estimated in the centimetreslice samples JF5–91b, -d, -e and -f, and phyllite sample
JF5–99 using the technique of Ferry (1992, 1994). Biotite,
chlorite, calcite and pyrrhotite modes were measured by
point counting in thin section and the modes of the other
six minerals were calculated from the whole-rock analyses
(Table 3) and given mineral compositions (Table 5) from
mass balance of SiO2, TiO2, Al2O3 and F
(MgO + FeO + Fe2O3 + MnO) (Table 6). Mineral
modes in the model protoliths were calculated for no
biotite and a pure albite plagioclase as observed in the
unreacted or barely reacted samples from locality 7
(Ferry, 1987). The thin slices from JF5–91 show that
the time-integrated fluid flux inferred from the reaction
progress varies by a factor of 50 on the centimetre scale
(Table 6) and correlates with rock composition, with
samples which contain the largest proportion of the pelitic
end-member exhibiting much higher amounts of reaction
(Fig. 14). Also, as previously inferred by Ferry (1994), the
samples with higher amounts of reaction have progressed
further across the divariant field of the assemblage to
slightly more water-rich fluid compositions. This is illustrated in Fig. 15 by a plot of the equilibrium constant
(ln K) for the equilibrium
KAl3Si3O10(OH)2 + 3CaMg(CO3)2 + 2SiO2=
muscovite
dolomite
quartz
KMg3AlSi3O10(OH)2 + CaAl2Si2O8 + 2CaCO3 + 4CO2
phlogopite
anorthite
calcite
against inferred time-integrated fluid flux. However, it
should be noted that the differences in ln K caused by
1% relative errors in the mineral analyses are small
although the values for JF5–91b and JF5–91d differ
at the 2r level (–1·79±0·12 vs –1·07±0·12). These
differences are consistent with the gradients in chemical
potential across the outcrop at locality 5 calculated by
Ferry (1979). If diffusive exchange between layers drove
reaction in the more pelitic layers, the minerals in the
more pelitic layers should have been in equilibrium with
more CO2-rich fluid compositions. Figure 15 shows that
the pelitic layers, with higher reaction progress, were, if
anything, in equilibrium with more water-rich fluids.
Explanations for the large differences in reaction progress and inferred fluid infiltration on the centimetre
scale include the following: (1) the infiltration and diffusion events occurred over different time scales, (2) reaction
began at higher temperatures or lower XCO2 in less pelitic
layers which therefore recorded a smaller fraction of the
fluid flow, or (3) compositional differences between layers
on the centimetre scale allowed diffusion of CO2–H2O
to drive more reaction in more pelitic layers, in which
reactions started at lower temperatures or higher fluid
XCO2 compositions.
A transient large porosity would enhance the fluid flux
far more than diffusional exchange. The time-integrated
fluid flux is directly proportional to the product of time
and permeability but permeability is a power-law function
of porosity (un/A). The fluid flux (xu) is therefore given
by
1507
JOURNAL OF PETROLOGY
VOLUME 38
NUMBER 11
NOVEMBER 1997
Table 6: Modal data in moles/litre
Sample:
JF5-91b
JF5-91d
JF5-91e
JF5-91f
JF5-99
Present assemblage
Muscovite
0·310
0·427
0·519
0·019
Biotite
0·0055
1·933
0·676
0·326
2·030
Chlorite
tr
0·066
0·021
0·007
tr
Calcite
15·17
Ankerite
2·247
Plagioclase
Quartz
9·155
13·79
2·470
18·43
0·055
1·170
1·484
0·585
0
0·902
2·157
1·525
0·848
2·963
7·238
tr
2·883
6·508
1·746
Rutile
0·068
0·071
0·076
0·044
0
Ilmenite
0
0
0
0
0·250
Pyrrhotite
0·028
0·475
0·113
0·056
0·056
4·595
Model protolith assemblage
Muscovite
0·315
2·329
1·156
0·343
Chlorite
0
0
0
0
Calcite
15·42
4·930
12·52
0·059
17·89
0
Ankerite
2·246
5·995
3·150
1·362
Siderite
0
0
0
0
1·242
3·231
Plagioclase
1·130
1·173
1·393
0·776
1·596
Quartz
6·317
1·772
2·813
6·386
3·439
Rutile
0·068
0·154
0·102
0·058
0·375
Pyrrhotite
0·028
0·475
0·113
0·056
0·056
Fluid production and time-integrated fluid flux from reaction progress
CO2∗
0·246
5·423
2·062
1·008
5·660
H2O∗
0·232
−0·304
0·067
0·075
0·333
NaCl=–HCl∗
RFlux (m3/m2)
0·463
150
−0·020
0·382
7030
2180
0·211
0
1009
22800
∗ +, moles/l produced; –, moles/l consumed.
Present rock calculated from point counting of biotite, chlorite, calcite and pyrrhotite with other minerals calculated
from whole-rock SiO2–TiO2–Al2O3–F (FeO+MgO+MnO)–CaO–K2O contents. Model protolith calculated from whole-rock
composition presuming no biotite and that plagioclase was pure albite. Protolith calculation for JF5-91d made by assuming
no chlorite and excluding K2O. tr, <0·05 modal %.
∂Pi un
∂z A
xu=
Fig. 14. Time-integrated fluid flux calculated from reaction progress
(see Ferry, 1994) for slices JF5–91b, -d, -e and -f plotted against the
fraction of phyllite end-member in the rock as calculated for Fig. 5a.
(11)
where n is in the range 2–3 for many porosity structures
(e.g. Dullien, 1979; Cheadle, 1989) and A is a constant.
Diffusional transport is proportional to the square-root
of the product of time and porosity [equation (1)]. Therefore, if a regime has a transient order of magnitude
increase of porosity for 1% of the time, the advective
flux might be increased by an order of magnitude (n=
3) but the diffusive exchange would only be increased
by 30%. Whether such porosity increases could preclude
diffusional exchange on the centimetre scale remains
open to question, as the diffusion distance for CO2 in a
porosity of 10–4 is ~1 cm in a year.
An alternative explanation is that the more pelitic
layers might start reacting at lower temperatures and
thus record higher time-integrated fluid fluxes. There are
1508
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
present mineral Fe/Mg ratios and this is expressed as
hypothetical pyrrhotite additional to that detected by
point counting (Table 7). The final assemblages in JF5–91
samples are related to their model protolith compositions
by reactions of the sort (e.g. for JF5–91d)
0·984musc + 2·327ank + 0·043rut + 0·891qz +
0·016H2O=1·0bio + 0·501anorth + 0·047alb +
(13)
1·955cc + 2·699CO2.
Fig. 15. Natural logarithm of equilibrium constant for reaction (10)
for present assemblage (Ε) and for model protolith assemblages at
biotite-in isograd assuming (1) no Na mobility (Φ) and (2) initial
plagioclase is An 2·3% (triangles), plotted against time-integrated fluid
flux calculated from reaction progress. Error bars represent 1r uncertainty propagated from 1% relative errors on microprobe analyses.
only small differences in fluid composition inferred for
equilibrium with the observed mineral assemblages across
all layers of JF5–91, and indeed all the carbonate band
samples. However, the more pelitic layers have high Fe/
Mg ratios and might start reacting at lower temperatures
than adjacent more carbonate-rich samples. The limited
divariance of the maximum phase assemblage is exhibited
by increases in Mg/Fe of biotite and chlorite, and anorthite content of plagioclase with increase in reaction
progress. This may be illustrated by calculating Fe/Mg
ratios of the hypothetical quartz–muscovite–
albite–anorthite–ankerite–calcite–rutile precursor assemblages. Mineral modes have been calculated from
bulk-rock compositions from Table 3 and mineral compositions in Table 5 in the composition space Si–
Ti–Al–F–Ca–Na–K where F is total Fe2+ + Mg + Mn
by solving the seven equations for six unknowns (amounts
of quartz, muscovite, albite, calcite and rutile in moles)
by a least-squares technique for given modes of biotite
and anorthite. Iron–magnesium distribution between
phases has been calculated to be consistent with the
average Fe–Mg distribution coefficients exhibited for the
minerals in Table 6, i.e. KDmusc–ank is given by
Fe mu Mgank
.
MgmuFeank
K Dmusc–ank=
(12)
The calculated mineral Fe/Mg ratios for the present
assemblage are consistently higher than analysed, which
may reflect the presence of iron-bearing opaque phases
in addition to those detected during point counting or
some systematic error in the calculation of mineral modes.
FeO has been reduced to give close matches to the
Ferry (e.g. 1994) observed that plagioclase compositions
at the biotite-in isograd are typically near pure albite (e.g.
An 1–3%) whereas plagioclase in the model precursor
assemblages calculated according to equation (13) would
contain An 10–30%. He argued that the fluid infiltration
associated with the biotite-forming reaction caused significant NaCl–HCl metasomatism such that the rocks
lost sodium. The model precursor assemblages have been
calculated for two assumptions: (1) at zero biotite content
and with anorthite content reduced according to the
stoichometry of reaction (13), and (2) at zero biotite
content and with the plagioclase set to An 2·3%. Protoliths
for the JF5–91 samples calculated as in method 1 exhibit
a range of initial plagioclase anorthite contents (Table 7)
with the implication that some anorthite was produced
by an additional reaction at lower grade. It should be
emphasized that all rocks with comparable bulk composition collected in the field area at the biotite-in isograd
have a near pure-albite plagioclase and there are no
other anorthite-producing reactions known in these rocks.
The probability of metasomatic compositional changes
associated with the fluid infiltration adds a further level
of complexity to interpretation of reaction progress in
these rocks.
The results are shown in Table 7 and Fe/(Mg + Fe)
of the first-formed biotite is plotted against calculated
time-integrated fluid flux in Fig. 16. Samples with higher
Fe/Mg ratios might start to react earlier than more
magnesium-rich samples and diffusional exchange during
this period could lower the necessary advective fluxes
and flux contrasts between layers. However, the variations
in Fe/Mg have relatively little effect on the width of the
divariant field for equilibrium (10), which is mainly
controlled by the change of plagioclase composition. This
is illustrated in Fig. 15, which plots the natural logarithm
of the equilibrium constant and the range of equilibrium
constants between biotite-in to present assemblage for
mineral compositions calculated as above (Table 7).
The third explanation for the variations of reaction
progress is that diffusive exchange of H2O and CO2
drives or impedes reaction progress in rocks within which
the fluid flux is relatively homogeneous. Layers with
more reactants might lose CO2 and gain H2O by diffusive
exchange with adjacent layers with less reaction in which
1509
JOURNAL OF PETROLOGY
VOLUME 38
NUMBER 11
NOVEMBER 1997
Table 7: Modal mineral proportions (moles/litre) and Fe/(Fe+Mg) ratios calculated for JF5-91
samples from whole-rock chemistry for protolith (no biotite) and observed (Table 6) biotite concentrations
JF5-91b
Sample
Mode
JF5-91d
JF5-91e
Fe/(Fe+Mg)
Mode
Fe/(Fe+Mg)
Mode
JF5-91f
Fe/(Fe+Mg)
Mode
Fe/(Fe+Mg)
Present assemblage
Musc
0·283
0·250
0·219
0·211
0·446
0·212
0·019
Biotite
0·0055
0·254
1·939
0·216
0·656
0·217
0·312
0·277
8·137
0·236
0·135
1·411
0·111
Calcite
15·00
13·54
0·237
18·36
Ankerite
2·156
1·328
0·503
Ab
0·876
1·264
1·214
0·613
An
0·324
1·697
0·762
0·328
Quartz
6·385
1·03
2·074
6·302
Rutile
0·067
0·077
0·077
0·044
Pyrrh
0·065
0·530
0·351
0·118
0·218
0·222
0·243
0·115
Protolith assemblage
Musc
0·288
0·250
1·995
0·320
1·048
0·274
0·325
0·267
Biotite
0
0·255
0
0·325
0
0·280
0
0·272
0·278
4·027
0·352
0·135
5·617
0·180
Calcite
14·98
12·02
2·804
0·304
0·150
17·60
Ankerite
2·168
1·209
Ab
0·875
1·106
1·165
0·592
An
0·322
0·728
0·426
0·171
Quartz
6·389
0·616
2·706
6·517
Rutile
0·68
0·150
0·101
0·058
Pyrrh
0·065
0·494
0·345
0·117
0·296
0·146
Pyrrhotite adjusted to give best fit Fe/Mg ratios in present assemblage.
Fig. 16. Fe/(Mg + Fe) molar ratio of protolith first-formed biotite
calculated from whole-rock composition (see text) plotted against timeintegrated fluid flux inferred from reaction progress.
infiltration has driven the divariant assemblage to equilibrium with a more water-rich fluid. In this case, ln K
should exhibit a negative correlation with reaction progress, but the opposite is observed (Fig. 15). The positive
correlation between reaction progress and fluid composition could be explained if the mineral assemblage
continues to re-equilibriate during changing pressures
and temperatures after the fluid infiltration, but the
precise mechanism which might result in a correlation
between fluid composition and reaction progress, other
than infiltration-driven reaction progress, is not clear.
With the precision of the information available at
present on the mineral equilibria it is not possible to
choose between the three possible hypotheses to explain
the rapid variations in apparent time-integrated fluid
flux, which are: (1) fluid flow was highly channelled in
the more permeable pelite-rich layers and took place in
short-lived high-flux events with limited diffusive
exchange, (2) fluid flow was pervasive and uniform, and
less pelitic horizons initiated the biotite-forming reactions
later at higher temperature or lower XCO2 and thus
recorded lower time-integrated fluid fluxes, and (3) reaction progress in the more pelitic layers was driven by
diffusion from adjacent unreacted more calcareous layers
in addition to layer-parallel infiltration.
1510
BICKLE et al.
FLUID FLOW, WATERVILLE LIMESTONE, MAINE
CONCLUSIONS
Boundary-layer profiles of d O, d C and Sr/ Sr isotope ratios across the margin of the Waterville limestone
at locality 5 all indicate advective displacements into the
limestone consistent with a time-integrated fluid flux of
3·2±1·4 m3/m2 (2r error). The profiles exhibit diffusive
broadening with a diffusion distance of 1·6±0·7 m for
d13C, 2·2±1·5 m for 87Sr/86Sr and 6·4±0·5 m for d18O.
The comparison of advective displacements and diffusion
distance for the three tracers is consistent with fluid being
relatively water rich (XCO2< 0·1) and saline (fluid Sr in
the range 75–400 p.p.m.). Small-scale Rb–Sr isochrons
indicate effective homogenization of Rb–Sr systematics
over distances of <1 m at ages within error of those
of the low-pressure Acadian metamorphic event and
associated granitic plutonism. This is consistent with the
fluid flow and diffusion event being associated with the
metamorphism. The small diffusion distances inferred
for oxygen and strontium isotopic homogenization at the
chlorite-grade locality 7 are consistent with this. As found
in previous studies, reaction progress varies by up to a
factor of 50 between adjacent limestone layers on the
centimetre scale and is up to a factor of 150 greater in
the adjacent phyllite. The metre-scale diffusion distances
and centimetre-scale variations in reaction progress might
result from short-lived high fluid flux events channelled
through more permeable, pelite-rich layers with the
diffusion taking place in lower-porosity conditions over
much longer time scales. Alternatively, the biotite-forming decarbonation reactions in the more pelite-rich layers
may have been initiated at lower temperatures or high
fluid XCO2 contents. If so, the differences in apparent
time-integrated fluid flux might result from the carbonaterich layers recording less fluid flow because they started
reacting later. This effect might have been accentuated
by diffusive exchange between the more CO2-rich pelitic
layers undergoing reactions and the adjacent unreacted
more carbonate-rich layers. The differences in predicted
mineral compositions are too small to test the relative
positions of the reactions in T– XCO2 space for the different
bulk compositions encountered. Irrespective of the precise
mechanisms for driving the biotite-producing decarbonation reaction, the extent of reaction progress in
the Waterville limestone requires substantial infiltration
of water-rich fluids. The isotopic boundary-layer profiles
indicate that the prevailing fluid was water rich and the
required fluid flux was probably largely layer parallel.
18
13
87
86
ACKNOWLEDGEMENTS
Judy Baker, Tim Holland and Katie Evans discussed
reaction progress in calc-silicate rocks. Ian Cartwright and
Alisdair Skelton provided thoughtful reviews. Research at
Cambridge on fluid movement in metamorphic rocks is
supported by NERC.
REFERENCES
Ague, J. J., 1994. Mass transfer during Barrovian metamorphism of
pelites, south–central Connecticut; I, Evidence for changes in composition and volume. American Journal of Science 294, 989–1057.
Armstrong, J. T., 1989. CITZAF: combined ZAF and phi–rho (Z)
electron beam corrections programs. Pasadena, CA: California Institute of Technology.
Baker, J. & Spiegelman, M., 1995. Modeling an infiltration-driven
geochemical front. Earth and Planetary Science Letters 136, 87–96.
Baumgartner, L. P. & Ferry, J. M., 1991. A model for coupled fluid-flow
and mixed-volatile mineral reactions with applications to regional
metamorphism. Contributions to Mineralogy and Petrology 106, 273–285.
Berner, R. A., Lasaga, A. C. & Garrels, R. M., 1983. The carbonate–
silicate geochemical cycle and its effect on atmospheric carbon
dioxide over the past 100 million years. American Journal of Science
283, 641–683.
Bickle, M. J., 1992. Transport mechanisms by fluid-flow in metamorphic
rocks: oxygen and strontium decoupling in the Trois Seigneurs
massif—a consequence of kinetic dispersion? American Journal of Science
292, 289–316.
Bickle, M. J., 1994. The role of metamorphic decarbonation reactions
in returning strontium to the silicate sediment mass. Nature 367,
699–704.
Bickle, M. J., 1996. Metamorphic decarbonation, silicate weathering
and the long-term carbon cycle. Terra Nova 8, 270–276.
Bickle, M. J. & Baker, J., 1990a. Advective–diffusive transport of isotopic
fronts: an example from Naxos, Greece. Earth and Planetary Science
Letters 97, 78–93.
Bickle, M. J. & Baker, J., 1990b. Migration of reaction and isotopic
fronts in infiltration zones: assessments of fluid flux in metamorphic
terrains. Earth and Planetary Science Letters 98, 1–13.
Bickle, M. J. & Chapman, H. J., 1990. Strontium and oxygen isotope
decoupling in the Hercynian Trois Seigneurs massif, Pyrenees:
evidence for fluid circulation in a brittle regime. Contributions to
Mineralogy and Petrology 104, 332–347.
Bickle, M. J. & McKenzie, D., 1987. The transport of heat and matter
by fluids during metamorphism. Contributions to Mineralogy and Petrology
95, 384–392.
Bickle, M. J., Wickham, S. M., Chapman, H. J. & Taylor, H. P., 1988.
A strontium, neodymium and oxygen isotope study of hydrothermal
metamorphism and crustal anatexis in the Trois Seigneurs massif,
Pyrenees, France. Contributions to Mineralogy and Petrology 100, 399–417.
Bickle, M. J., Chapman, H. J., Wickham, S. M. & Peters, M. T., 1995.
Strontium and oxygen isotope profiles across marble–silicate contacts, Lizzies Basin, East Humbolt Range, Nevada: constraints on
metamorphic fluid flow. Contributions to Mineralogy and Petrology 121,
400–413.
Borthwick, J. & Harmon, R. S., 1982. A note regarding ClF3 as an
alternative to BrF5 for oxygen isotope analysis. Geochimica et Cosmochimica Acta 46, 1665–1668.
Carslaw, H. S. & Jaeger, J. C., 1959. Conduction of Heat in Solids. Oxford:
Oxford University Press, 510 pp.
Cheadle, M. J., 1989. Properties of texturally equilibrated two-phase
aggregates. Ph.D. Thesis, University of Cambridge.
Clayton, R. N. & Kieffer, S. W., 1991. Oxygen isotope thermometer
calibrations. In: Taylor, H. P., Jr, O’Neil, J. R. & Kaplan, I. R. (eds)
Stable Isotope Geochemistry: a Tribute to Samuel Epstein. Geochemical Society
Special Publications 3, 3–10.
1511
JOURNAL OF PETROLOGY
VOLUME 38
Clayton, R. N. & Mayeda, T. K., 1963. The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for
isotopic analysis. Geochimica et Cosmochimica Acta 27, 43–52.
Dallmeyer, R. D. & Van Breeman, O., 1981. Rb–Sr whole-rock and
40
Ar/39Ar mineral ages of the Togus and Hallowell quartz monzonite
and Three Mile Pond granodiorite plutons, south–central Maine:
their bearing on post-Acadian cooling history. Contributions to Mineralogy and Petrology 78, 61–73.
de Laeter, J. R. & Abercrombie, I. D., 1970. Mass spectrometric isotope
dilution analyses of rubidium and strontium in standard rocks. Earth
and Planetary Science Letters 9, 327–330.
Dullien, F. A. L., 1979. Porous Media Fluid Transport and Pore Structure.
New York: Academic Press, 396 pp.
Eiler, J. M., Baumgartner, L. P. & Valley, J. W., 1992. Intercrystalline
stable isotope diffusion—a fast grain boundary model. Contributions
to Mineralogy and Petrology 112, 543–557.
Etheridge, M. A., Wall, V. J. & Vernon, R. H., 1983. The role of the
fluid phase during regional metamorphism and deformation. Journal
of Metamorphic Geology 1, 205–226.
Eugster, H. P., Albee, A. L., Bence, A. E., Thompson, J. B. & Waldbaum, D. R., 1972. The two-phase region and excess mixing properties of paragonite–muscovite crystalline solutions. Journal of Petrology
13, 147–179.
Ferry, J. M., 1976. P, T, f CO2 and f H2O during metamorphism of
calcareous sediments in the Waterville–Vassalboro area, south–
central Maine. Contributions to Mineralogy and Petrology 57, 119–143.
Ferry, J. M., 1979. A map of chemical potential differences within an
outcrop. American Mineralogist 64, 966–985.
Ferry, J. M., 1980. A case study of the amount and distribution of heat
and fluid during metamorphism. Contributions to Mineralogy and Petrology
71, 373–385.
Ferry, J. M., 1986. Infiltration of aqueous fluid and high fluid:rock
ratios during greenschist facies metamorphism: a reply. Journal of
Petrology 27, 695–714.
Ferry, J. M., 1987. Metamorphic hydrogeology at 13-km depth and
400–550°C. American Mineralogist 72, 39–58.
Ferry, J. M., 1988. Contrasting mechanisms of fluid flow through
adjacent stratigraphic units during regional metamorphism, south–
central Maine, USA. Contributions to Mineralogy and Petrology 98, 1–12.
Ferry, J. M., 1992. Regional metamorphism of the Waits River Formation, Eastern Vermont: delineation of a new type of giant metamorphic hydrothermal system. Journal of Petrology 33, 45–94.
Ferry, J. M., 1994. Overview of the petrological record of fluid flow
during regional metamorphism in northern New England. American
Journal of Science 294, 905–988.
Ganor, J., Mathews, A. & Paldor, N., 1989. Constraints on effective
diffusivity during oxygen isotope exchange at a marble–schist contact,
Sifnos (Cyclades), Greece. Earth and Planetary Science Letters 94, 208–
216.
Garlick, G. D. & Epstein, S., 1967. Oxygen isotope ratios in coexisting
minerals of regionally metamorphosed rocks. Geochimica et Cosmochimica Acta 31, 181–214.
Goodge, J. W. & Holdaway, M. J., 1995. Rock-buffered fluid evolution
of metapelites and quartzites in the Picuris Range, Northern New
Mexico. Journal of Petrology 36, 1229–1250.
Helgeson, H. C., 1967. Solution chemistry and metamorphism. In:
Abelson, P. H. (ed.) Researches in Geochemistry. New York: John Wiley,
pp. 362–404.
Hoisch, T. D., 1987. Heat transfer by fluids during late Cretaceous
metamorphism in the Big Maria Mountains, southeastern California.
Geological Society of America Bulletin 98, 549–553.
NUMBER 11
NOVEMBER 1997
Holland, T. J. B. & Powell, R., 1990. An enlarged and updated
internally consistent thermodynamic dataset with uncertainties and
correlations:
the
system
K2O–Na2O–CaO–MgO–MnO–
FeO–Fe2O3–Al2O3–SiO2–C–H2–O2. Journal of Metamorphic Geology 8,
89–124.
Holland, T. J. B. & Powell, R., 1992. Plagioclase feldspars: activity–
composition relations based upon Darken’s quadratic formalism and
Landau theory. American Mineralogist 77, 53–61.
Kerrick, D. M. & Caldeira, K., 1993. Paleoatmospheric consequences
of CO2 released during early Cenozoic regional metamorphism in
the Tethyan orogen. Chemical Geology 108, 201–230.
Lasaga, A. C. & Rye, D. M., 1993. Fluid flow and chemical reaction
kinetics in metamorphic systems. American Journal of Science 293,
361–404.
Matthews, A., 1994. Oxygen isotope geothermometers for metamorphic
rocks. Journal of Metamorphic Geology 12, 211–219.
Osberg, P. H., 1968. Stratigraphy, structural geology, and metamorphism of the Waterville–Vassalboro area, Maine. Maine Geological
Survey Bulletin 20, 64 pp.
Osberg, P. H., 1988. Geologic relations within the shale–wacke
sequence in South–Central Maine. In: Studies in Maine Geology. Maine
Geological Survey, pp. 51–73.
Palmer, M. R., 1992. Controls over the chloride concentration of
submarine vent fluids: evidence from Sr/Ca and 87Sr/86Sr ratios.
Earth and Planetary Science Letters 109, 37–46.
Press, W. H., Flannery, B. P., Teulolsky, S. A. & Vetterling, W. T.,
1986. Numerical Recipes. Cambridge: Cambridge University Press, 818
pp.
Rumble, D., Ferry, J. M., Hoering, T. C. & Boucot, A. J., 1991. Carbon
and oxygen isotope geochemistry of chlorite-zone rocks of the
Waterville limestone, Maine, USA. American Mineralogist 76, 857–866.
Rye, R. O., Schuiling, R. D., Rye, D. M. & Jansen, J. B. H., 1976.
Carbon, hydrogen, and oxygen isotope studies of the regional
metamorphic complex at Naxos, Greece. Geochimica et Cosmochimica
Acta 40, 1031–1049.
Sharp, Z. D. & Kirschner, D. L., 1993. Quartz–calcite oxygen isotope
thermometry: an empirical calibration. Geological Society of America,
Abstracts with Programs.
Skelton, A. D. L., Graham, C. M. & Bickle, M. J., 1995. Lithological
and structural controls of regional 3-D fluid flow patterns during
greenschist facies metamorphism of the Dalradian of the SW Scottish
Highlands. Journal of Petrology 36, 563–586.
Skelton, A. D. L., Bickle, M. J. & Graham, C. M., 1997a. Fluid-flux
and reaction rate from advective–diffusive carbonation of mafic sill
margins in the Dalradian, southwest Scottish Highlands. Earth and
Planetary Science Letters 146, 527–539.
Skelton, A. D. L., Graham, C. M., Bickle, M. J. & Cole, C., 1997b.
Lithological, structural and deformation controls on fluid flow during
regional metamorphism. In: Holness, M. B. (ed.) Deformation Enhanced
Fluid Flow in the Earth’s Crust and Mantle. London: Chapman and Hall,
in press.
Walther, J. V. & Orville, P. M., 1982. Volatile production and transport
in regional metamorphism. Contributions to Mineralogy and Petrology 79,
252–257.
Wood, B. J. & Graham, C. M., 1986. Infiltration of aqueous fluid and
high fluid:rock ratios during greenschist facies metamorphism: a
discussion. Journal of Petrology 27, 751–761.
Yardley, B. W. D. & Lloyd, G. E., 1995. Why metasomatic fronts are
really metasomatic sides. Geology 23, 53–56.
York, D., 1969. Least squares fitting of a straight line with correlated
errors. Earth and Planetary Science Letters 5, 320–324.
1512