ELSEVIER
Earth and Planetary Science Letters 169 (1999) 227–244
www.elsevier.com/locate/epsl
Cracking at the magma–hydrothermal transition: evidence from the
Troodos Ophiolite, Cyprus
K.M. Gillis a,Ł , M.D. Roberts b
a
School of Earth and Ocean Sciences, P.O. Box 3055, University of Victoria, Victoria, BC V8W 3P6, Canada
b Economic Geology Research Unit, James Cook University, Townsville, Qld 4811, Australia
Received 10 August 1998; revised version received 1 March 1999; accepted 12 March 1999
Abstract
The nature of the magma–hydrothermal transition in oceanic hydrothermal systems is poorly understood, in part
because the geological relations in this critical region have rarely been observed in modern ocean crust. Detailed mapping
was conducted in the Troodos Ophiolite, Cyprus, where a gabbronorite sequence intrudes the sheeted dyke complex, which
is truncated at its base by a thin contact aureole composed of massive hornfels. Geothermometric data for hornblende
and pyroxene hornfels show that hydrated sheeted dykes were recrystallized at amphibolite to granulite facies conditions
(778–986ºC). Quartz diorite veins and apophyses, and monomineralic amphibole veins cross-cut the contact aureole and
show no preferred age relationships. Geothermometric data indicate that quartz diorite was injected at 817–919ºC and
that fractures were filled with amphibole at 575–750ºC. Phase relations of quartz-hosted, halite-bearing fluid inclusions in
quartz diorite veins constrain minimum entrapment temperatures of 225–520ºC (average 402ºC) and minimum pressures
that span lithostatic and hydrostatic conditions. We believe that these characteristics are indicative of a conductive boundary
layer that separates an active hydrothermal system from the heat source that drives it. Field and petrological data indicate
that transient fracturing caused oscillations in temperature and pressure conditions within the conductive boundary layer,
and mixing of hydrothermal and magmatic fluids at the magma–hydrothermal interface. Cross-cutting relations between
magmatic and hydrothermal vein networks show that fracturing occurred prior to the cessation of magmatic activity.
We explore plausible models for the causes and consequences of fracturing that consider the role of dyke injection,
thermoelastic stresses, and volatile build-up.  1999 Elsevier Science B.V. All rights reserved.
Keywords: Troodos Ophiolite; geothermal systems; contact metamorphism; aureoles
1. Introduction
There is little direct evidence concerning the
nature of the magma–hydrothermal transition in
oceanic hydrothermal systems. Vent fluid compositions, in combination with experimental and theoŁ Corresponding
[email protected]
author. Fax: C1 250 721 6200; E-mail:
retical studies, constrain conditions within the rootzones of hydrothermal cells [1,2]. In this region,
hydrothermal fluid compositions become fixed by
reaction with the surrounding rocks at low fluid=rock
ratios, near supercritical temperatures, and hydrostatic pressures (500 bar), prior to their ascent to
the seafloor (see review by [1]). The constancy of
fluid chemistries, in particular soluble elements such
as Li and B, on decadal time scales (i.e., the time pe-
0012-821X/99/$ – see front matter  1999 Elsevier Science B.V. All rights reserved.
PII: S 0 0 1 2 - 8 2 1 X ( 9 9 ) 0 0 0 8 7 - 4
228
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
riod vents have been sampled) requires that circulating fluids are continuously in contact with fresh rock
[1], either by the migration of hydrothermal rootzones or injection of fresh material by periodic dyking. Input of magmatic volatiles into hydrothermal
systems is well documented [3], however, the link
between the flux specific volatile species and magma
chamber processes remains poorly constrained.
Conceptual and mathematical models place the
magma–hydrothermal transition at or near the sheeted
dyke–plutonic boundary [4]. A critical aspect of these
models is the presence of a conductive boundary layer
that separates a vigorously convecting hydrothermal
cell from the heat source that drives it [4–7]. These
layers are modeled as dynamic features whose thickness and position in the crust varies during the life
span of a hydrothermal system, depending on the
balance between the heat supplied by the convecting
magma and heat lost by conduction across the layer
[6]. The thickness of the conductive layer varies with
magmatic heat flux in that active hydrothermal systems with high thermal discharge (¾103 MW) have
very thin layers (meters) whereas those with lower
thermal discharge (¾10 MW) have layers on the order of a few hundred meters [7,8]. Models require
that conductive boundary layers remain thin in order
to maintain high heat flux hydrothermal systems on
decadal time scales, by the downward migration of a
cracking front into hot rock, or unknown magmatic or
tectonic processes [7,8].
Ophiolites offer a complementary venue for
studying oceanic hydrothermal processes. Indeed,
Cyprus-type volcanogenic massive sulphide deposits
are among the best known ancient analogues for
modern, mid-ocean ridge sulphide deposits (e.g.,
TAG [9]). Although most key elements of ophiolite-hosted hydrothermal systems are similar to those
at modern mid-ocean ridges, important differences
exist in regard to the extent and conditions of alteration within sheeted dyke complexes. Ophiolite
sheeted dyke complexes are more uniformly and
pervasively altered, at higher fluid=rock ratios, than
modern sheeted dyke complexes and contain large
(up to 8 km2 ) zones of epidosite where dykes have
been extensively metasomatized and recrystallized.
Epidosites represent either the reaction zones or base
of hydrothermal upflow zones [10,11] and are rare in
mid-ocean ridge rock collections.
In this paper we present new geological and petrological constraints about the nature of the magma–
hydrothermal transition in oceanic hydrothermal systems. We selected the Troodos Ophiolite for study
as a wealth of information has been accumulated
since the 1960s concerning its magmatic, tectonic,
and hydrothermal evolution (see [12]). Our mapping
builds on the early work of Bear [13] and Malpas and Brace [14] in an area where intrusion of
a sequence of gabbroic rocks into the base of the
sheeted dyke complex produced a thin contact aureole. We interpret the contact aureole as a preserved
conductive boundary layer which isolated an active
hydrothermal cell from its magmatic heat source and
document evidence for extreme oscillations in P –T
conditions.
2. Geologic relations
2.1. Background
The Troodos Ophiolite, located on the island of
Cyprus in the eastern Mediterranean, lies at the
northwestern end of a belt of Cretaceous ophiolites
that are discontinuously exposed along the margin of
the Arabian plate. The ophiolite is Turonian in age
[15] and is thought to have formed in a supra-subduction zone setting within the Tethys Ocean [16].
Obduction of the Troodos oceanic crust involved a
90º anticlockwise rotation between the Maastrichtian
and early Eocene and episodic uplift that culminated
in the Early to Mid-Quaternary [12]. Emplacementrelated metamorphism was minimal and is restricted
to late-stage fracture and fault-infilling [17].
Estimates for spreading rate of the Troodos ocean
crust based on structural features and paleoseafloor
topography span the range of possibilities (e.g.,
[18,19]). What is more relevant to this study is
that field relationships suggest temporal changes in
magma supply during crustal construction [20]. Field
relations between the sheeted dyke complex and plutonic sequence, and between different plutonic bodies clearly show that the Troodos crust was built by
multiple magma chambers [21,22]. The most compelling evidence of multiple intrusive relations is the
common occurrence of xenoliths of one lithology in
a host gabbro of slightly different composition [23].
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
Complex relations between the sheeted dykes and
gabbros also support models for polyphase magmatism. Sheeted dykes are commonly intruded by high
level gabbroic plutons, which may be later intruded
by dyke swarms, and, in a few locations, dykes arise
directly from the pluton [21,24]. The relief along this
transition is ½500 m as dyke swarms locally intrude
the plutonics from depth and plutons locally intrude
to high levels.
2.2. Platanistasa Window
The Platanistasa Window is located along the
northern flank of the ophiolite, between the villages of Alona, Platanistasa, and Polystipos (Fig. 1).
A gabbronorite sequence intrudes the sheeted dyke
complex, which is truncated at its base by a thin
contact aureole composed of massive, fine grained
hornfels. The relief along this contact is <50 m except where it is locally down-dropped <100 m to the
east along high angle normal faults.
229
The sheeted dykes strike N to NW, dip steeply to
the east (¾60º), and range in thickness from ¾0.5
to >4 m. There is no systematic increase in dyke
thickness towards the plutonics, and dykes rooted directly in the plutonic sequence, as described by Allen
[24], were not observed. The contact aureole, which
separates the basal sheeted dyke complex from the
subjacent gabbronorite sequence, is locally cut by a
network of quartz diorite veins and apophyses, and
monomineralic amphibole veins (Fig. 2).
The plutonic sequence is dominantly composed
of massive to layered gabbronorite, olivine gabbronorite, and wehrlite, with lesser amounts of gabbro, magnetite gabbro, diorite, quartz diorite, and
plagiogranite. Layered gabbronorites extend up to
the base of the sheeted dykes along most of the
contact and are locally infiltrated by leucocratic vein
networks or dykes (Fig. 3, left). The proportion of
leucocratic veins in outcrop systematically increases
from <10% to 30% as the contact is approached.
Varitextured amphibole gabbro occurs within the
Fig. 1. Geologic map of the Platanistasa Window. The contact aureole is centered along the igneous boundary between the sheeted
dykes and layered series. Topographic relief exposes a vertical section (400–600 m) across the sheeted dyke–plutonic transition along
two NE–SW ridges, such that from north to south, sheeted dykes ! plutonics ! sheeted dykes crop out. Inset shows the generalized
geology of the Troodos Ophiolite. Contour interval is 200 feet.
230
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
Fig. 2. (A) Outcrop photograph and (B) sketch showing a sharp intrusive contact where gabbro intrudes into pyroxene hornfels. Note that
the contact is cross-cut by an amphibole vein network. (C) Outcrop photograph and (D) sketch showing quartz diorite and amphibole
veins cross-cutting very fine grained hornfels.
layered sequence as massive outcrops, dykes, and
veins, and has a wide range in texture and grain size
on a deci-meter scale. Plagiogranite bodies locally
intrude into the dykes at the same structural level as
the contact aureole and have intrusive or gradational
boundaries with the gabbronorite sequence (Fig. 1).
Outcrops range in composition from diorite to quartz
diorite to plagiogranite and contain partially resorbed
xenoliths of hornfelsic basalt stoped from the base of
the sheeted dyke complex. Elsewhere, gabbro, gabbronorite, magnetite gabbro, and varitextured gabbro
with complex intrusive relationships form the roofzone assemblage. Isolated dykes which locally intrude the layered gabbronorite sequence strike NNE
and are east-dipping (¾60º).
Sheeted dykes and lavas exposed north of the
Platanistasa Window are relatively unfaulted and un-
tilted, and lie between the Solea and Mitsero grabens
[18]. The prevalence of an intrusive sheeted dyke–
plutonic boundary and absence of significant dyke
rotation are indicative of formation during a phase of
magmatic extension. By contrast, the extreme crustal
attenuation in the Solea graben records a period of
waning magmatism or amagmatic extension [20].
3. Metamorphic evolution
In this section mineral assemblages and compositions, and textural features of key lithologies are
presented in order to determine the protolith of the
hornfels and constrain the thermal conditions in the
vicinity of the contact aureole. Mineral compositions
for key lithologies are summarized in Table 1.
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
231
Fig. 3. (Left) Schematic diagram showing typical geological relationships along the sheeted dyke–plutonic transition. Sheeted dykes
(vertical lines) are truncated at their base by the gabbronorite sequence (light gray). As the contact aureole is approached, the abundance
of felsic melt impregnations (solid squares) within the gabbronorites increases (up to 30%). The contact aureole is cross-cut by quartz
diorite (white lines C solid squares) and amphibole (white lines) veins; these veins show no preferred age relationships. (Right)
Thermal conditions in the vicinity of the contact aureole. See text for details. Abbreviations: SD D sheeted dykes; HH D hornblende
hornfels; PH D pyroxene hornfels; QD D quartz diorite; GS D gabbronorite series; gm D groundmass; 1 D mineral assemblage; 2 D
plagioclase–amphibole geothermometer; 3 D orthopyroxene–clinopyroxene geothermometer; 4 D lava chemistry.
Table 1
Summary of mineral compositions a
Lithology
Sheeted dykes
Hornblende hornfels
Pyroxene hornfels
Quartz diorite
Gabbronorite sequence
CY4 sheeted dykes
CY4 upper gabbros
Plagioclase
Clinopyroxene Orthopyroxene Amphibole
Mg# b
Mg# b
(AlIV ) c
Mg# b
Primary
Secondary
An40–80 (avg. An65 )
An35–80 (peaks at An44–48 , An56–60 )
An65–95
An34–64
An85–90 ; An53–60
An49–90 (avg. An60–70 )
An80–95
An0–15 d
–
An5–15 d
0.64–0.78
An40–55 e ; An0–15 0.56–0.74
–
An0–15 d
0.5–0.9
–
0.64–0.74
–
0.74–0.84
–
–
0.56–0.74
–
0.70–0.85
0.59–0.66
0.68–0.78
0.11–1.04
0.27–1.15
0.14–1.30
0.28–1.25
0.11–1.52
–
–
0.28–0.68
0.48–0.82
0.50–0.76
0.36–0.90
0.34–0.96
–
–
a Data were collected using JEOL 733 and 8900 microprobes at the Massachusetts Institute of Technology and University of Alberta,
respectively; standard Bence–Albee corrections were used.
b Mg# D Mg2C =(Mg2C C Fe2C ).
c Amphiboles were recalculated on the basis of 23 anhydrous oxygens following the normalization of Robinson et al.
d Adjacent to microfractures or patches within primary grain.
e Narrow rims adjacent to amphibole.
232
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
3.1. Petrologic features
The lowermost 200–300 m of the sheeted dyke
complex is pervasively altered to greenschist facies
mineral assemblages. Clinopyroxene is completely
replaced by fibrous to granular amphibole, plagioclase is partially albitized, and interstitial zones are
replaced by amphibole š quartz š chlorite š epidote. Where present, chlorite and quartz are minor
phases (<5 modal%) and epidote occurs in trace
amounts. Epidositized patches, where igneous minerals are completely replaced by epidote C quartz
š chlorite š magnetite, were observed within a few
isolated dykes. Larger scale epidosite bodies have
been mapped to the NW [10] and NE of the Platanistasa Window.
The contact aureole that separates the sheeted
dyke complex and gabbronorite sequence is composed of massive, very fine grained outcrops of
hornblende and=or pyroxene hornfels. Where both
types of hornfels were identified, the hornblende
hornfels formed further from the gabbronorite sequence than the pyroxene hornfels. Quartz diorite
veins and apophyses, and amphibole veins cut the
contact aureole. Amphibole vein networks with random orientations cross-cut some quartz diorite veins
and are cut by other quartz diorite veins, suggesting
that they formed during multiple fracturing events.
Quartz diorite veins contain plagioclase C quartz C
amphibole š clinopyroxene š Fe oxides š zircon
š apatite and are typically coarser grained than the
host hornfels.
Hornblende hornfels are characterized by the assemblage granular amphibole C clinopyroxene C
plagioclase C quartz C relict fibrous amphibole C
relict clinopyroxene C magnetite C ilmenite. The
extent of recrystallization varies from <20 to 100%
over distances of mms to cms, with most samples
showing <50% recrystallization. Zones with relict
minerals have diabasic to intersertal textures that resemble the hydrothermally altered sheeted dykes,
however, there is a notable absence of chlorite
and epidote in the hornfels. Relict clinopyroxene
is almost (>80%) completely replaced by fibrous
amphibole whereas recrystallized clinopyroxene is
finer grained, granoblastic, and unaltered. Granular
amphibole has an interstitial to poikiolitic texture
(Fig. 4A) or rims clots of fibrous amphibole that
pseudomorphically replace clinopyroxene (Fig. 4B),
suggesting that it formed by the recrystallization of
fibrous amphibole. Plagioclase commonly retains its
igneous habit but may have sutured grain boundaries
and subgrain development in areas where clinopyroxene is recrystallized.
Pyroxene hornfels are characterized by the assemblage clinopyroxene C orthopyroxene C plagioclase C granular amphibole C magnetite C ilmenite š quartz š relict fibrous amphibole. The degree of recrystallization ranges from ¾20 to 100%,
with most samples showing >50% recrystallization.
Relict zones display diabasic textures and contain fibrous amphibole, plagioclase, and quartz; no igneous
clinopyroxene is preserved. Recrystallized zones
contain discrete grains of granoblastic clinopyroxene
and orthopyroxene (Fig. 4C) or aggregates of grains
whose form mimics relict pyroxene grains (Fig. 4D).
Pyroxene aggregates contain >75% clinopyroxene,
minor orthopyroxene (<10%), and trace quartz
(<5%), and are locally mantled by poikiolitic brown,
granular amphibole. Plagioclase typically retains its
igneous habit and is coarser grained than associated
granoblastic pyroxene (Fig. 4C). Less commonly,
plagioclase has a granoblastic texture and equivalent
grain size to pyroxene (Fig. 4D) or exhibits subgrain
development.
The gabbronorite sequence is variably altered to
greenschist to amphibolite facies mineral assemblages. The extent of alteration is less uniform and
pervasive than the sheeted dykes. Pervasive replacement of igneous phases is generally localized along
vein margins and areas infiltrated by felsic magmas.
3.1.1. Protolith of the contact aureole
Petrological features indicate that the contact
aureole is dominantly composed of recrystallized
sheeted dykes. The hornblende hornfels have diabasic to intersertal textures, and primary mineral compositions are most similar to the sheeted dykes. The
pyroxene hornfels have features of both the sheeted
dykes and gabbros, suggesting that they represent
both recrystallized dykes and gabbros.
3.2. Temperature constraints
Peak temperatures are constrained by mineral
compositions and assemblages, in conjunction with
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
233
Fig. 4. Photomicrographs showing typical textures and mineral assemblages in the hornfels. (A) Hornblende hornfels with granular,
brown amphibole (ga) poikiolitically enclosing plagioclase (pl) laths (sample KG93090). (B) Hornblende hornfels showing fibrous green
amphibole (fa) replacing clinopyroxene (c) and rimmed with granular amphibole (ga) (sample KG93056). (C) Pyroxene hornfels with
very fine grained, granular aggregates of clinopyroxene and orthopyroxene (c-o); note that plagioclase (pl) and the pyroxene aggregates
have similar grain size (sample KG93181). (D) Pyroxene hornfels showing a very fined grained granoblastic assemblage of plagioclase,
orthopyroxene, clinopyroxene, magnetite, and ilmenite (sample KG93171). Field of view is 1.5 mm.
experimental studies [25–27] and using two-pyroxene [28] and plagioclase–amphibole [29] thermometers (Table 2).
The long-lived thermal conditions prevalent in
the vicinity of the contact aureole are recorded by
the metamorphic mineralogy (Fig. 3B). Hydrothermal mineral assemblages within the basal sheeted
dykes indicate temperatures between 400 and 550ºC
whereas granular, groundmass amphibole formed
at higher temperatures (600–770ºC). Why higher
temperatures are recorded by groundmass amphibole rather than the other hydrothermal minerals is
not clear. One possibility is that granular amphibole formed during peak metamorphic conditions
and that the remaining metamorphic minerals record
conditions during the waning stages of hydrothermal
convection. Alternatively, these mineral assemblages
may record temperature fluctuations during the life
span of a hydrothermal system. Transient increases
in temperature may lead to the formation of granular amphibole, by the breakdown of greenschist
facies minerals such as chlorite and actinolite (see
discussion below).
Within the contact aureole, hornblende and pyroxene hornfels record recrystallization temperatures
of 778–904ºC and 940–986ºC, respectively. Quartz
diorite veins cross-cutting the aureole crystallized
at magmatic temperatures (817–919ºC), whereas the
amphibole veins sealed fractures at 575–750ºC. Temperatures abruptly increased across the upper boundary with the sheeted dykes, from either ¾450 or
700ºC to 830ºC. Thermal conditions across the lower
boundary are more difficult to constrain as magmatic
temperatures for the gabbronorite sequence have not
234
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
Table 2
Calculated metamorphic temperatures (ºC) a
Sample #
Groundmass amphibole b
Sheeted dykes
KG93011
KG93036
KG93038
KG93143
KG93188
KG93067
689–770 (2)
732 (1)
612–734 (4)
710–718 (2)
636 (1)
–
Hornblende hornfels
KG93056
850–875
KG93090
549–904
KG93098
892 (1)
KG93125
752–862
KG93215
779 (1)
KG93217
774–802
(2)
(2)
(4)
(2)
Pyroxene hornfels
KG92056
823 (1)
KG92091
806–897 (5)
KG93023
649–940 (2)
KG93079
–
KG93171
–
KG93181
730–950 (2)
Gabbronorites
KG92074
KG93030
KG93204
KG93205
KG93209
KG93210
KG93213
673–835
691–845
652–937
775–819
653–778
717–838
778–839
(8)
(4)
(5)
(7)
(7)
(3)
(2)
Quartz diorite vein amphibole b
–
–
–
–
–
–
Vein amphibole b
493 (1)
–
–
–
–
615–775 (8)
–
–
–
–
–
–
556–929 (6)
864–898 (3)
823–865 (2)
–
–
–
810–847 (14)
–
–
–
–
575–657 (3)
–
659–698 (5)
609–661 (6)
–
–
–
–
–
–
–
–
–
–
773 (1)
–
–
–
–
–
Groundmass pyroxene c
–
–
–
–
–
–
–
–
–
–
–
–
969–972 (2)
872–909 (3)
953–1020 (4)
940–986 (2)
939–973 (3)
910–972 (4)
–
–
–
–
–
–
–
a Temperature
range for each sample; number of analyses is shown in parentheses.
Plagioclase–amphibole thermometers of Holland and Blundy [25].
c Two-pyroxene thermometer of Andersen et al. [24].
b
been calculated. Glass and phenocryst compositions
for the lava sequence indicate crystallization temperatures between 1225 and 1000ºC [30]. Temperatures
clearly increased across this boundary from ¾960ºC
to super-solidus conditions.
4. Microthermometry
Microthermometric fluid inclusion data were collected for five quartz diorite veins which cross-cut
the hornblende and pyroxene hornfels (Table 3).
Variations in room-temperature phase ratios, homogenization temperatures, and salinity suggest multiple
populations of fluid inclusions with variable chemical and thermal histories [31]. The following section
focuses on quartz-hosted, halite-bearing fluid inclusions in order to constrain the evolution and trapping
conditions of high-salinity fluids within the contact
aureole. Phase relations for halite-bearing fluid inclusions, which largely homogenize by halite dissolution, indicate trapping of a magmatic fluid at lithostatic conditions and relatively low temperatures.
4.1. Methods
Microthermometric determinations were carried
out on a Fluid Inc. heating and freezing petrographic
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
235
Table 3
Summary of fluid inclusion data
Sample
Lith. a
Type
n
Th b
(min–max)
AVG d
Tm c
(min–max)
NaCl e (wt%)
(min–max)
KG92091
PH
1
2
14
10
170–220
179–395
194
249
–
275–520
3.4–20.2
38.6–62.5
KG93090
HH
1
2
23
8
167–325
220–350
230
270
–
225–440
2.5–5.3
33.2–52.0
KG93098
HH
1
15
221–345
302
KG93125
HH
2
1
19
6
190–376
230–330
245
295
–
215–403
1.5–7.1
32.7–47.8
KG93127
HH
1
2
21
13
206–367
198–373
307
–
372–478
2.3–3.3
44.5–56.8
–
2.6–4.9
a PH
D pyroxene hornfels, HH D hornblende hornfels.
Th D temperature of vapor bubble homogenization (ºC).
c T D temperature of halite dissolution (ºC).
m
d AVG D average value of n inclusions.
e Salinity calculated as wt% NaCl equivalent.
b
stage following methods outlined in Roedder [32].
Heating and freezing data were collected from individual fluid inclusions with a range in size from 10
to 20 µm. Consistency of phase ratios and homogenization temperatures among fluid inclusions within
secondary arrays was used to ensure representative
data. Measurement accuracy is estimated at š0.1ºC
for final ice-melting determinations (š0.2 wt% NaCl
equivalent) and š5ºC for liquid–vapor homogenization (Th ) and halite melting (Tm ) events (š1.0 wt%
NaCl eq.). Salinity calculations were performed using the fluid inclusion software package MacFlincor
(version 0.77), with the equations of Brown and
Lamb [33] for the H2 O–NaCl–[KCl] system. Fluid
inclusions are modeled in the H2 O–NaCl system and
salinity is reported as wt% NaCl equivalent.
4.2. Fluid inclusion types
Fluid inclusions hosted within quartz are categorized into two types: (1) primary and secondary, lowsalinity, liquid-dominated inclusions and (2) primary
and secondary, halite-bearing, liquid-dominated inclusions. Distinction between primary versus secondary inclusion genesis was equivocal in some
samples due to the abundance of inclusions.
Type 1 low-salinity fluid inclusions range between
1.5 and 8.6 wt% NaCl eq., and exhibit homogeniza-
tion to the liquid at temperatures between 167 and
376ºC. Salinity values fall within the observed salinity range of fluids emanating from active hydrothermal vents (70% below to 200% above seawater
values) [2]. Several rare liquid-dominated inclusions
yielded salinities of ¾20 wt% NaCl eq.
Type 2 halite-bearing, liquid-dominated inclusions appear to be primary and secondary in nature and are identified by the presence of one or
more cubic halite daughter minerals. Halite-bearing
inclusions yield three modes of homogenization behavior whereby the vapor bubble disappears at (1)
temperatures less than halite melting temperatures
(Th < Tm ), (2) temperatures equal to halite melting temperatures (Th D Tm ), and (3) temperatures
greater than halite melting temperatures (Th > Tm ).
Most inclusions display mode (1) behavior (92%)
with vapor bubble homogenization temperatures of
179–395ºC and fluid salinities of 36–63 wt% NaCl
(Tm D 277–520ºC) (Fig. 5). It should be noted that
inclusions of 25–35 wt% NaCl may be under-represented due to the common failure of fluid inclusions
within this salinity range to nucleate a halite crystal [34]. Many halite-bearing inclusions contain a
reddish and=or black opaque daughter mineral that
are tentatively identified as hematite and a sulphide
phase, respectively.
Quartz-hosted fluid inclusions from fault gouge
236
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
Fig. 5. Temperatures of halite dissolution (Tm ) and vapor bubble disappearance (Th ) for halite-bearing fluid inclusions. Black
diamonds represent fluid inclusions hosted within quartz diorite veins which cross-cut the pyroxene and hornblende hornfels;
clear circles represent fluid inclusions hosted within evolved plutonic rocks at the base of the contact aureole; the shaded field
represents previously published data for fluid inclusions hosted
within bodies of plagiogranite, epidosite and pegmatitic gabbro
located proximal to the sheeted dyke–plutonic transition within
the Troodos Ophiolite (from [42]).
samples and evolved plutonic rocks from the field
area display similar populations with several key
differences [31]. Fluid inclusions within fault gouge
samples yield low salinities and are mostly liquiddominated. In contrast, fluid inclusions from evolved
plutonic rocks contain both low-salinity and halitebearing inclusions, as well as vapor-rich inclusions.
4.3. Conditions of fluid entrapment
The minimum temperature of trapping for halitebearing inclusions which homogenize by halite dissolution (Th < Tm ) or by simultaneous vapor bubble
disappearance and halite dissolution (Th D Tm ) is
given by the temperature of halite dissolution [32].
As such, the trapping of high-salinity fluid (36–
63 wt% NaCl) within the quartz diorite veins occurred at minimum temperatures of 225–520ºC (avg.
402ºC). The lack of fluid inclusions which homogenize to the vapor phase, and the majority of halitebearing inclusions which homogenize by halite dissolution (Th < Tm ) indicate that the brines were
not trapped as an immiscible fluid [32]. This ob-
Fig. 6. (A) Temperature-pressure phase diagram for 40 wt%
NaCl fluids in the H2 O–NaCl system. The curve L C V C H
represents the pressure and temperature conditions whereby liquid, vapor and halite coexist in equilibrium ([58] and references
therein); the critical curve separates the one-phase liquid field
from the two-phase liquid–vapor field (subdivided into salinitydependent sub-fields) [59]; the dashed line represents the halite
liquidus for a 40 wt% NaCl fluid [34]. The hashed region and
the dot-dashed line represent possible P –T condition of trapping for fluid inclusions of 40 wt% NaCl which exhibit mode
1 (Th < Tm ) and mode 2 (Th D Tm ) homogenization behavior,
respectively. Trapping conditions for inclusions which display
mode 3 homogenization behavior (Th > Tm ) exist at temperatures greater than the line Th D Tm . Depth is modelled as a
function of lithostatic pressure beneath the seafloor (see text).
servation, in conjunction with consistent room temperature phase ratios, suggests that a single phase,
homogeneous fluid was trapped.
Isochoric projections were constructed to further
constrain possible P –T conditions of fluid entrapment (Fig. 6). Using the equations of Bodnar and
Vityk [34] isochores for the halite-bearing fluid inclusions hosted within the quartz diorite veins were
projected from the three-phase curve (L C V C
H) at their temperatures of vapor bubble homogenization. It should be noted that these equations are
only valid for fluid inclusions of 0–40 wt% NaCl
with homogenization temperatures between 50 and
700ºC at pressures below 6 kbar [34]. Isochores for
halite-bearing inclusions with salinities greater that
40 wt% NaCl must be approximated using the slopes
for inclusions of 40 wt% NaCl (Bodnar, pers. commun., 1998) because experimental data for fluids
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
with higher salinities are lacking. In addition, the
true slope of isochores in the liquid C halite field
have a slightly more positive slope than isochores
in the one-phase liquid field [35]. Therefore, these
projections only represent a first approximation of
P –T trapping conditions for the halite-bearing inclusions.
Based on the H2 O–NaCl model system, halitebearing fluid inclusions which homogenized by
halite dissolution (Th < Tm ) have salinities that
range from 36–63 wt.% NaCl. These inclusions
were trapped in P –T space between their respective liquidus and the line which represents trapping
conditions for inclusions which exhibit simultaneous
vapor bubble disappearance and halite dissolution
(Th D Tm ) (Fig. 6) [35]. Intersection of an isochore
for a halite-bearing fluid inclusion that homogenized
by halite dissolution (Th < Tm ) and its respective liquidus represents the minimum pressure of trapping
for that inclusion [32]. Homogenization temperature
data for halite-bearing fluid inclusions, modeled as
40 wt% NaCl fluids (see above), describe isochores
which intersect the halite liquidus at pressures between 900 and 1200 bar (Fig. 6). Approximate isochores for inclusions with salinities greater than 40
wt% NaCl support minimum trapping pressures of
¾1000 bar. Although all isochores for inclusions
which homogenize by halite dissolution (Th < Tm )
indicate a minimum pressure of trapping, inclusions
for which the temperature of vapor bubble disappearance is significantly lower than halite dissolution
record higher minimum pressures of trapping [36].
To place these minimum pressures in a geologic
context, a model of hydrostatic and lithostatic pressure gradients with depth was constructed. Assuming
Cretaceous ocean depths of 2000 m, a rock density
of 2950 kg=m3 , and a water density of 1000 kg=m3 ,
lithostatic pressures of 900–1200 bar correspond to
depths of 2275–3175 m below the seafloor, whereas
the same pressures under hydrostatic conditions correspond to depths >7200 m below the seafloor. Field
relations indicate that the paleo-depth of the contact aureole was 2000–2500 m below the seafloor.
Based on available constraints these minimum trapping pressures suggest conditions of fluid entrapment
at lithostatic pressures or greater, whereas trapping
under hydrostatic conditions places the contact aureole at unrealistic depths.
237
5. Discussion
The prevalence of an intrusive sheeted dyke–
plutonic boundary in the Platanistasa Window is indicative of formation during a phase of magmatic
extension. Abrupt truncation of the basal dykes at
the contact aureole implies that the pluton intruded
and assimilated the roots of the basal dykes. Age relationships for veins cross-cutting the aureole show
that it was multiply fractured and subjected to significant oscillations in pressure and temperature (see
below). We believe that the dynamic nature of the
contact aureole is best explained by formation during a phase of active hydrothermal circulation and
magmatic crustal construction.
Many features of the contact aureole exposed in
the Platanistasa Window are indicative of a conductive boundary layer. It is situated at the top of a
magma chamber, at the base of the sheeted dykes,
in an area that has undergone very little tectonic
extension. The contact aureole is thin (¾10–30 m)
and records steep thermal gradients across its boundaries. The overlying sheeted dyke complex is more
uniformly and pervasively altered than the subjacent
gabbronorite sequence. In the following section we
explore the evolution of fluid compositions and P –
T conditions in the vicinity of the contact aureole
and develop a dynamic model that requires transient
fracturing and mixing of hydrothermal and magmatic
fluids at the magma–hydrothermal interface.
5.1. Evolution of fluid compositions
Numerous fluid inclusion studies have documented the presence of high-salinity fluids within
gabbroic rocks recovered from the modern ocean
crust and its ophiolitic analogues (see [37]). Phase
separation of seawater has been cited as a possible mechanism for the salinity and phase variations observed within hydrothermal vent fluids and
deep-seated fluid inclusions (see [38]). However, few
oceanic hydrothermal fluid inclusion studies have
unequivocally identified fluids trapped as immiscible liquid=vapor pairs [37,39,40]. Vapor–brine density contrasts and rock fracture configurations have
been suggested as mechanisms to explain high-salinity inclusions in the absence of vapors [40,41]. In
contrast, it has been argued that the physical prop-
238
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
erties of seawater coupled with rapid healing rates
of quartz micro-fractures would inhibit the migration
of seawater further into the two-phase field, thereby
limiting an end-member liquid phase to less than 30
wt% NaCl and hindering segregation of cogenetic
liquids and vapors [42]. Consequently, Kelley [37]
has argued that halite-bearing fluid inclusions from
evolved plutonic rocks which homogenize by halite
dissolution indicate trapping of high-salinity fluids
exsolved from late stage melts in the absence of a
vapor phase.
Halite-bearing fluid inclusions which homogenize
by halite dissolution are trapped in the absence of a
vapor phase [32]. This condition requires that a highsalinity fluid of seawater or magmatic origin must be
cooled substantially prior to trapping (Fig. 6) [35].
Although high-salinity, magmatic fluids can be exsolved into either the one- or two-phase field, phase
separation of high-salinity fluids generates a low
mass fraction of vapor to liquid [36]. The pervasive
absence of cogenetic, low-salinity vapor condensates
and the microthermometric constraints provided by
the halite-bearing fluid inclusions trapped within the
quartz diorite veins strongly suggests a magmatic
origin for the high-salinity fluids. Although no temporal relationships could be determined, the halitebearing fluid inclusions are hosted among numerous populations of secondary, near-seawater salinity
fluid inclusions. Mixing of hydrothermal fluids with
higher temperature magmatic fluids is a possible
mechanism for cooling the high-salinity magmatic
fluids. The upper minimum pressure of trapping suggests that the high-salinity fluids were trapped, after
cooling, near to or greater than lithostatic conditions.
We conclude that taken together, the origin and
condition under which the high-salinity fluids were
trapped, reflect extreme fluctuations in temperature
and pressure within the conductive boundary layer.
5.2. Oscillations in P –T conditions
Cracking events occur within contact aureoles in
response to large physical and chemical gradients
which vary in time and space (see [43]). In the
Platanistasa Window, the steady-state thermal conditions within the contact aureole are constrained
by the hornfels which document high temperatures
(778–986ºC). Fluid inclusion data and features of
the aureole indicate that lithostatic conditions dominated. Phase relations for halite-bearing fluid inclusions constrain minimum trapping pressures of 0.9–
1.2 kbar at lithostatic conditions. Recrystallization
temperatures for the hornfels are higher than plausible temperatures for the brittle–plastic transition
in oceanic crust, placing the aureole at lithostatic
pressures [44]. Constraints for the brittle–plastic
transition are based on dry diabase flow laws (see
discussion in Hirth et al. [44]) and petrologic evidence that incipient cracking within oceanic gabbros
occurs at ½700ºC (e.g., [45]).
Constraints from the vein networks show that
steady-state temperature and pressure conditions
were periodically perturbed within the conductive
boundary layer. Amphibole vein networks and fluid
inclusions trapped along microfractures hosted in
quartz diorite veins record fracture sealing at 575–
750ºC and 225–520ºC, respectively. Quartz diorite
veins and apophyses intruded at magmatic temperatures (817–919ºC). The macroscopic vein networks
suggest that fracturing reduced pressures from lithostatic to hydrostatic conditions within the aureole. As
these vein types show no preferred age relationships,
we conclude that P –T conditions oscillated during
the life span of the conductive boundary layer.
Theoretical studies and field observations at modern mid-ocean ridges provide a basis for understanding the mechanisms that created the fracture
networks within the Platanistasa conductive boundary layer. Dyke intrusion is an extensional event
which brings magma to the seafloor from a magma
chamber residing beneath the conductive boundary
layer. Dyke injection occurs within a narrow zone
(tens of meters) (see [46]), at intervals of years to
centuries. Theoretical constraints suggest that high
temperatures within the conductive boundary layer
cause a build-up in thermoelastic stresses, compressive along the cool side of the boundary, tensile
along the hot [47]. As a dyke penetrates this compressive layer, fractures will rapidly propagate into
the underlying layer that is under tension. Theoretical constraints also predict that mechanical stresses
within a sheeted dyke complex caused by dyke injection would close fractures parallel to the plane
of a dyke and open fractures oriented perpendicular to the dyke [48]. These horizontal tensile and
compressive stresses would extend well beyond the
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
dyke injection zone, affecting the entire axial rise at
a fast-spreading ridge (i.e., up to 1.5 km from dyke
margins) [48]. Over several years, thermal contraction fractures would open and mechanical stresses
would decrease to ambient levels. Thus, dyke injection may trigger a fracturing event within the
conductive boundary layer, and will close and open
fractures within the overlying sheeted dyke complex
over a time span of several years.
Thermoelastic stresses may also create permeability at the boundary between hot rock and cool
seawater [5,6]. This is the basis of Lister’s conceptual model which invokes a cracking front that
migrates down into the crust, extracting heat from
the hot rock through which it is penetrating. Cracking occurs at temperatures below the rigidus, which
for basaltic rocks is probably ½700ºC [44,45]. The
cracking front is commonly thought of in terms of
the waning stage of a magma body as it moves out
of the zone of dyke injection and off-axis. It may
also develop during the life span of a magma body in
response to fluctuations in magma supply [49].
Other mechanisms for fracturing involve magmatic and=or metamorphic fluids. Fluids exsolved
by the crystallization of a melt act to increase the
total pressure within the magma chamber until the
maximum tensile strength of the overlying layer is
exceeded. In shallow environments, water saturation
and volatile exsolution is generally not achieved until late in the crystallization history (>90%) [50].
This process can be episodic, with subsequent cracking events requiring less overpressure to fracture the
overlying rock or to reopen pre-existing fractures.
The early prograde path of contact metamorphism
is characterized by thermally induced pore fluid expansion and metamorphic devolatilization reactions
which can produce pore fluid pressures in excess
of the ambient lithostatic load (¾3–25 MPa), which
are sufficient to hydrofracture the host rocks [51].
Other processes, however, such as crack healing and
pressure solution result in an overall permeability
structure dominated by intergranular fluid flow [43].
5.3. Evolution of the Platanistasa Window
conductive boundary layer
A model that integrates geological and petrological constraints from the Platanistasa area is illus-
239
Fig. 7. Stages in the evolution of the Platanistasa conductive
boundary layer. (A) During steady-state conditions a contact
aureole (dark shading) separates sheeted dykes from a magma
chamber below. A sharp thermal gradient exists at the top of
the contact aureole where the conductive boundary layer (dashed
line) separates hydrothermal fluids circulating under hydrostatic
pressures from pore fluids at or above lithostatic conditions.
(B) Episodic dyking generates fractures within the contact aureole, allowing penetration of hydrothermal fluids to deeper
levels above a thinned conductive boundary layer. Sudden release in pressure promotes the expulsion of magmatic volatiles
and allows for mixing of magmatic and hydrothenual fluids. (C)
Fractures are subsequently sealed by mineral precipitation which
effectively shifts the conductive boundary layer upwards. The
new thermal structure results in recrystallization and assimilation of the injected dyke. (D) Cessation of magmatic activity
is marked by the downward migration of thermal contraction
fractures such that the conductive boundary layer subsides into
the gabbronorite sequence.
trated in Fig. 7. A steady-state hydrothermal system
has been established above a magma chamber. The
conductive boundary layer coincides with the contact
aureole which is composed of previously hydrated
sheeted dykes that are undergoing recrystallization to
hornblende and pyroxene hornfels. Lithostatic pressures (¾1 kbar) and high temperatures (>750ºC)
are prevalent within the conductive boundary layer.
There is a steep thermal gradient between the conductive boundary layer and overlying hydrothermal
system where hydrothermal fluids are convecting at
temperatures ½350ºC. The subjacent magma cham-
240
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
ber is at super-solidus conditions and is undergoing
complex crystallization involving the assimilation of
previously stoped, hydrated dyke roots and magmatic
differentiation. Fracturing disrupts these steady-state
conditions and facilitates the migration of magmatic
and hydrothermal fluids into the conductive boundary layer from below and above, respectively.
As outlined in the previous section, we must rely
on theoretical studies and observations from modern mid-ocean ridges to incorporate fracturing into
our model for the Platanistasa conductive boundary
layer. Field and petrological data document oscillations in P –T conditions. Cross-cutting relations
between the magmatic and hydrothermal vein networks show that fracturing occurred prior to the
cessation of magmatic activity. We cannot constrain,
however, if the preserved features of the conductive boundary layer formed within an axial zone of
spreading or off-axis during the waning stages of
magmatic activity.
If we consider that the conductive boundary layer
evolved within an axial zone, steady-state conditions
(Fig. 7A) are perturbed by the injection of a dyke
into the brittle crust creating or re-opening fractures within the conductive boundary layer due to
thermoelastic and mechanical stresses (Fig. 7B). Enhanced permeability causes the overlying hydrothermal system to deepen and the conductive boundary
layer to thin which in turn increases the heat flux of
the hydrothermal system [7,52]. Within this newly
fractured zone, temperatures decrease and pressures
oscillate from lithostatic to hydrostatic conditions
(¾1 kbar to ¾400 bar). Sudden release in pressure
promotes the expulsion of magmatic fluids into the
overlying hydrothermal system, as well as metamorphic pore fluids generated by dehydration reactions
within the contact aureole, causing the complex mixing of magmatic, metamorphic, and hydrothermal
fluids (Fig. 7B) [43]. Magmatic and metamorphic
fluids cool by mixing with cooler, hydrothermal fluids. A rapid drop in pressure, from lithostatic to
hydrostatic conditions, promotes crystallization of
the quartz diorite veins [50]. At the same time, hydrothermal fluids from the overlying system infiltrate
the newly fractured zone and are heated by mixing with hotter magmatic and metamorphic fluids
and=or by conduction from the wallrock. Temperature increase beyond the quartz solubility maxima
leads to quartz precipitation [53] and the sealing
of microfractures within the quartz diorite veins at
pressures between hydrostatic and lithostatic. The
abundance of amphibole veins within the aureole
suggests that increasing hydrothermal fluid temperatures also promotes the precipitation of amphibole
(e.g., [27]); however, we are not familiar with any
theoretical calculations of amphibole solubility.
As the permeability is reduced by mineral precipitation, the conductive boundary layer thickens
upward until the system returns to steady-state conditions whereby the magmatic heat input is balanced
by hydrothermal heat extraction (Fig. 7C). By this
time, lithostatic pressures and high temperatures are
restored within the conductive layer, causing the root
of the recently injected dyke to be recrystallized
within the contact aureole and assimilated within the
roof-zone of the magma chamber. This completes a
typical cycle (Fig. 7B,C).
The time scale of a typical cycle is not well constrained but is likely to be similar to the frequency of
dike intrusion (<10–100 years). Seismic monitoring
of mid-ocean ridges suggests that dyke propagation
lasts in the order of days [54]. Thermal effects in the
wall rock adjacent to an intruding dyke last hours to
days [55]. Mechanical stresses are restored to ambient values within several years [48]. Thus, fracture
propagation associated with dyking is short-lived.
Rates of fracture sealing by mineral precipitation
are assumed to be rapid at high temperatures, with
the time scale being dependent on parameters such
as fracture geometry, fluid compositions, reaction
kinetics, and flow velocities [56]. At the P –T conditions documented for the Platanistasa conductive
boundary layer, microfracture sealing rates for quartz
in equilibrium with saline solutions have time scales
of hours to days [57]. At conditions above the quartz
saturation maxima, quartz precipitation would reduce the porosity by a few volume percent on a
time scale of decades [56]. Thus, it is probable that
the permeability created by dyking is significantly
reduced prior to the next dyking event.
The thickness of the conductive boundary layer
is, in part, linked to the time interval between dyking
events. During extended periods of quiescence, the
conductive boundary layer may thicken up into the
sheeted dyke complex [49], beyond the upper boundary of the contact aureole, promoting dehydration re-
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
actions that lead to the formation of amphibole. The
extent of recrystallization would be limited by the
permeability of the basal dykes and the longevity of
the thermal excursion. Within the Platanistasa Window, heterogeneous distribution of high temperature
amphibole within sheeted dykes altered to greenschist facies mineral assemblages is consistent with
transient upward migration of higher temperature
conditions.
In summary, steady-state conditions within the
conductive boundary layer developed above a
magma chamber within the axial zone of spreading (Fig. 7C). Dyke injection causes the conductive boundary layer to fracture, which facilitates the
migration and mixing of magmatic and hydrothermal fluids into the conductive boundary layer and
the sealing of fractures by mineral precipitation
(Fig. 7B). As steady-state conditions are restored
(Fig. 7C), a typical cycle is completed. The conductive boundary layer undergoes multiple cycles prior
to moving out of the axial zone. The time scale of
each cycle is comparable to the frequency of dyking.
An equally plausible model is that the Platanistasa
conductive boundary layer solely records conditions
during the waning stages of magmatic activity. In this
off-axis setting, exsolution of magmatic fluids within
the roof-zone of the magma chamber increases the
total pressure beyond the maximum tensile strength
of the thinned boundary layer and induces hydrofracturing [50]. As outlined above, fracturing would
cause a change from lithostatic to hydrostatic conditions, and mixing of magmatic, metamorphic, and
seawater-derived hydrothermal fluids within the conductive boundary layer. More than one fracturing
event is required to produce the distribution of magmatic veins cutting hydrothermal veins, and vice
versa.
The final event for both our proposed models,
which consider the features formed either within
an axial zone or off-axis, is the solidification of
the magma chamber (Fig. 7D). At this time, the
conductive boundary layer migrates down into the
gabbronorite sequence, followed by a zone of incipient cracking — the cracking front of Lister [5,6].
Cracking is initiated at conditions equivalent to the
brittle–ductile transition, and sealed at temperatures
of 575–750ºC.
In this section, we have attempted to reconstruct
241
the causes and consequences of fracturing within
the conductive boundary layer exposed in the Platanistasa Window. We have outlined two plausible
models that are consistent with our geological and
petrological data, theoretical considerations, and observations from modern mid-ocean ridges. Clearly,
other models that consider the timing of dyking,
volatile build-up, and cracking fronts in different
ways are equally plausible.
6. Summary
The prevalence of an intrusive sheeted dyke–
plutonic boundary in the Platanistasa Window is indicative of formation during a phase of magmatic
extension. Abrupt truncation of the basal dykes at
a contact aureole implies that the pluton intruded
and assimilated the roots of the basal dykes. Many
features of the contact aureole are indicative of
a conductive boundary layer that separates an active hydrothermal system from the heat source that
drives it. The contact aureole formed within the
basal sheeted dykes and uppermost gabbros due to
the intrusion of a layered gabbronorite sequence.
Mineralogical characteristics of the hornfels indicate
that the basal dykes were hydrated prior to recrystallization, by previous or on-going hydrothermal
processes, and that recrystallization occurred at amphibolite to granulite facies conditions (778–986ºC)
and lithostatic pressures. Field and petrological data
document oscillations in P –T conditions. Cross-cutting relations between magmatic and hydrothermal
vein networks show that fracturing occurred prior to
the cessation of magmatic activity.
We speculate that the fracture networks formed
either within the axial zone of spreading or during
the waning stages of magmatic activity. As a starting point, the hydrothermal system is considered to
be at steady-state conditions (i.e., between fracturing events) and the conductive boundary layer is
located at or near the top of the contact aureole.
If the preserved features of the conductive boundary layer form within the axial zone, dyke injection
causes fracturing within the conductive boundary
layer, causing it to thin and migrate toward the base
of the contact aureole. Fracture propagation promotes the mixing of magmatic, metamorphic, and
242
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
hydrothermal fluids, and leads to the deposition of
quartz diorite and amphibole veins. Subsequent fracture sealing results in the resumption of steady-state
conditions. The conductive boundary layer is subjected to multiple fracturing events prior to moving
out of the axial zone. If the preserved features of
the conductive boundary layer record conditions during the waning stages of magmatic activity, volatile
build-up induces hydrofracturing, which again facilitates the mixing of magmatic, metamorphic, and hydrothermal fluids and the formation of quartz diorite
and amphibole veins. Once a magma chamber solidifies off-axis, oscillations in temperature and pressure
cease, and the conductive boundary layer migrates
downward into the plutonic sequence. We recognize
that we present two of many plausible models for
the causes and consequences of fracturing within the
conductive boundary layer.
[3]
[4]
[5]
[6]
[7]
[8]
[9]
Acknowledgements
We are grateful to C. Xenophontos and the Cyprus
Geological Survey Department for their on-going
support of geological research programs. We thank
the Hellenic Mining Company for logistical support
in the field. The focus of this project evolved through
many lively discussions in the field with J. Cann,
J. Malpas, C. Xenophontos, K. Deveau, S. Agar,
L. Marquez, and K. Klitgord. We acknowledge N.
Chatterjee and R. Luth for their assistance with
microprobe analyses and K. Sapp for her assistance
with many aspects of the project during its early
stages. We also thank W. Wilcox and an anonymous
reviewer for their thorough reviews. This research
was supported by NSF (EAR-9118791) and NSERC
research grants. [CL]
[10]
[11]
[12]
[13]
[14]
[15]
References
[16]
[1] P.J. Saccocia, K. Ding, M.E. Berndt, J.S. Seewald, W.E.
Seyfried, Experimental and theoretical perspectives on
crustal alteration at mid-ocean ridges, in: Lentz, D.R.
(Ed.), Alteration and Alteration Processes Associated with
Ore-forming Systems, Vol. 11, Geol. Assoc. Canada Short
Course Notes, St. John’s, 1994, pp. 403–432.
[2] K. Von Damm, Controls on the chemistry and temporal
variability of seafloor hydrothermal fluids, in: Humphris,
[17]
[18]
[19]
S.E., Zierenberg, R.A., Mullineaux, L.S., Thomson, R.E.
(Eds.), Seafloor Hydrothermal Systems, Geophys. Monogr.
Ser. 91, Am. Geophys. Union, Washington, DC, 1995, pp.
222–247.
D. Kadko, J. Baross, J. Alt, The magnitude and global
implications of hydrothermal flux, in: Humphris, S.E.,
Zierenberg, R.A., Mullineaux, L.S., Thomson, R.E. (Eds.),
Seafloor Hydrothermal Systems, Geophys. Monogr. Ser. 91,
Am. Geophys. Union, Washington, DC, 1995, pp. 446–466.
R.P. Lowell, P.A. Rona, R.P. Von Herzen, Seafloor hydrothermal systems, J. Geophys. Res. 100 (1995) 327–352.
C.R.B. Lister, On the penetration of water into hot rock,
Geophys. J. R. Astron. Soc. 39 (1974) 465–509.
C.R.B. Lister, On the intermittency and crystallization
mechanisms of sub-seafloor magma chambers, Geophys.
J. R. Astron. Soc. 73 (1983) 351–366.
C.R.B. Lister, Heat transfer between magmas and hydrothermal systems, or, six lemmas in search of a theorem,
Geophys. J. Int. 120 (1995) 45–59.
R.P. Lowell, L.N. Germanovich, Dike injection and the
formation of megaplumes at ocean ridges, Science 267
(1995) 1804–1807.
S.E. Humphris, P.M. Herzig, D.J. Miller et al., The internal
structure of an active sea-floor massive sulphide deposit,
Nature 377 (1995) 713–716.
C.J. Richardson, J.R. Cann, H.G. Richards, J.G. Cowan,
Metal-depleted root zones of the Troodos ore-forming hydrothermal systems, Cyprus, Earth Planet. Sci. Lett. 84
(1987) 243–253.
P. Schiffman, B.M. Smith, R.J. Varga, E.M. Moores, Geometry, conditions and timing of off-axis hydrothermal metamorphism and ore-deposition in the Solea graben, Nature
325 (1987) 423–425.
A. Robertson, C. Xenophontos, Development of concepts
concerning the Troodos ophiolite and adjacent units in
Cyprus, in: Prichard, H.M., Alabaster, T., Harris, N.B.W.,
Neary, C.R. (Eds.), Magmatic Processes and Plate Tectonics, Geol. Soc. Spec. Publ. 76 (1993) 85–119.
L.M. Bear, The geology and mineral resources of the
Akaki–Lythrodondha area, Government of Cyprus, Nicosia,
1960, 122 pp.
J. Malpas, T. Brace, The geology of Pano Amiandos–
Palekhori area, Cyprus, Geological Survey Department,
Nicosia, 1987.
S.B. Mukasa, J.N. Ludden, Uranium-lead isotopic ages of
plagiogranites from the Troodos ophiolite, Cyprus and their
tectonic significance, Geology 15 (1987) 825–828.
J.A. Pearce, S.J. Lippard, S. Roberts, Characteristics and
tectonic significance of supra-subduction zone ophiolites,
Geol. Soc. London 16 (1984) 77–94.
K.M. Gillis, P.T. Robinson, Patterns and processes of alteration in the lavas and dykes of the Troodos Ophiolite,
Cyprus, J. Geophys. Res. 95 (1990) 21523–21548.
R.J. Varga, E.M. Moores, Spreading structure of the Troodos ophiolite, Cyprus, Geology 13 (1985) 846–850.
S. Allerton, F.J. Vine, Spreading structure of the Troodos
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
[20]
[21]
[22]
[23]
[24]
[25]
[26]
[27]
[28]
[29]
[30]
[31]
[32]
[33]
[34]
ophiolite, Cyprus: some paleomagnetic constraints, Geology 15 (1987) 593–597.
Varga, E.M. Moores, Intermittent magmatic spreading and
tectonic extension in the Troodos Ophiolite: implications
for exploration of black smoker-type ore deposits, in: Malpas, J., Moores, E.M., Panayiotou, A., Xenophontos, C.
(Eds.), Ophiolites: Oceanic Crustal Analogues, Cyprus Geological Survey Department, Nicosia, 1990, pp. 53–64.
J. Malpas, T. Brace, S.M. Dunsworth, Structural and petrologic relationships of the CY-4 drill hole of the Cyprus
Crustal Study Group, in: Gibson, I.L., Malpas, J., Robinson, P.T., Xenophontos, C. (Eds.), Cyprus Crustal Study
Project: Initial Report, Hole CY-4, Geological Survey of
Canada Paper 88-9, Geological Survey of Canada, Ottawa,
1989, pp. 39–68.
E.M. Moores, F.J. Vine, The Troodos Massif, Cyprus and
other ophiolites as oceanic crust: evaluation and implications, Phil. Trans. R. Soc. London A 268 (1971) 443–466.
J. Malpas, Crustal accretionary processes in the Troodos
ophiolite, Cyprus: evidence from field mapping and deep
crustal drilling, in: Malpas, J., Moores, E.M., Panayiotou,
A., Xenophontos, C. (Eds.), Ophiolites: Ocean Crustal Analogues, Geological Survey Department, Nicosia, 1990, pp.
65–74.
C.R. Allen, The petrology of a portion of the Troodos plutonic complex, Cyprus, Ph.D. thesis, Cambridge University,
1975.
J.G. Liou, S. Kuniyoshi, K. Ito, Experimental studies of the
phase relations between greenschist and amphiobolite in a
basaltic system, Am. J. Sci. 274 (1974) 613–632.
M.J. Apted, J.G. Liou, Phase relations among greenschist,
epidote–amphibolite, and amphibolite in a basaltic system,
Am. J. Sci. 283A (1983) 328–354.
F.S. Spear, An experimental study of hornblende stability
and compositional variability in amphibolite, Am. J. Sci.
281 (1981) 697–734.
D.L. Andersen, D.H. Lindsley, P.M. Davidson, Quilf: A
Pascal program to assess equilibria among Fe–Mg–Mn–Ti
oxides, pyroxenes, olivine and quartz, Comput. Geosci. 19
(1993) 1333–1350.
T. Holland, J. Blundy, Non-ideal interactions in calcic amphiboles and their bearing on amphibole–plagioclase thermometry, Contrib. Mineral. Petrol. 116 (1994) 433–447.
P. Thy, C. Xenophontos, Crystallization orders and phase
chemistry of glassy lavas from the pillow sequences, Troodos Ophiolite, Cyprus, J. Petrol. 32 (1991) 403–428.
M.D. Roberts, The role of magmatic fluids and phase
separation at the sheeted dyke–plutonic transition: a fluid
inclusion study from the Troodos Ophiolite, Cyprus, B.Sc.
thesis, University of Victoria, 1998.
E. Roedder, Fluid inclusions, Mineralogical Society of
America, Washington, DC, 1984, 644 pp.
P.E. Brown, W.M. Lamb, P –V –T properties of fluids in
the system H2 O š CO2 š NaCl: new graphical presentations and implications for fluid inclusions studies, Geochim.
Cosmochim. Acta 53 (1989) 1209–1221.
R.J. Bodnar, M.O. Vityk, Interpretation of microthermo-
[35]
[36]
[37]
[38]
[39]
[40]
[41]
[42]
[43]
[44]
[45]
[46]
[47]
[48]
243
metric data for H2 O–NaCl fluid inclusions, in: De Vivo, B.,
Frezzotti, M.L. (Eds.), Fluid Inclusions in Minerals: Methods and Applications, Virginia Polytechnic Institute and
State University, Pontignana–Siena, 1994, pp. 117–130.
R.J. Bodnar, Synthetic fluid inclusions: XII. The system
H2 O–NaCl. Experimental determination of the halite liquidus and isochores for a 40 wt% NaCl solution, Geochim.
Cosmochim. Acta 58 (1994) 1053–1063.
J.S. Cline, D.A. Vanko, Magmatically generated saline
brines related to molybdenum at Questa, New Mexico,
USA, in: Thompson, J.F.H. (Ed.), Magmas, Fluids, and Ore
Deposits, Mineral. Assoc. Canada Short Course Ser. 23,
Mineralogical Association of Canada, Nepean, 1995, pp.
153–174.
D.S. Kelley, Fluid evolution in slow-spreading environments, in: Karson, J.A., Cannat, M., Miller, D.J., Elthon,
D. (Eds.), Proc. ODP, Sci. Results, Vol. 153, Ocean Drill.
Program, College Station, TX, 1997, pp. 399–415.
P. Nehlig, Interactions between magma chambers and hydrothermal systems: oceanic and ophiolitic constraints, J.
Geophys. Res. 98 (1993) 19621–19633.
D.S. Kelley, J.R. Delaney, Two-phase separation and fracturing in mid-ocean ridge gabbros at temperatures greater
than 700ºC, Earth Planet. Sci. Lett. 83 (1987) 53–66.
D.A. Vanko, R.J. Bodnar, S.M. Sterner, Synthetic fluid
inclusions: VIII. Vapor-saturated halite solubility in part
of the system NaCl–CaCl2 –H2 O, with application to fluid
inclusions from oceanic hydrothermal systems, Geochim.
Cosmochim. Acta 52 (1988) 2451–2456.
M.S. Goldfarb, J.R. Delaney, Response of two-phase fluids
to fracture configurations within submarine hydrothermal
systems, J. Geophys. Res. 93 (1988) 4585–4594.
D.S. Kelley, P.T. Robinson, J.G. Malpas, Processes of brine
generation and circulation in the oceanic crust: fluid inclusion evidence from the Troodos Ophiolite, Cyprus, J.
Geophys. Res. 97 (1992) 9307–9322.
R.B. Hanson, The hydrodynamics of contact metamorphism, Geol. Soc. Am. Bull. 107 (1995) 595–611.
G. Hirth, J. Escartin, J. Lin, The rheology of the lower
oceanic crust: implications for lithospheric deformation at
mid-ocean ridges, in: Buck, W.R., Delaney, P.T., Karson,
J.A., Lagabrielle, Y. (Eds.), Faulting and Magmatism at
Mid-Ocean Ridges, Geophys. Monogr. Ser. 106, American
Geological Union, Washington, DC, 1998.
C. Manning, P.E. Weston, K.I. Mahon, Rapid high temperature metamorphism of the East Pacific Rise gabbros from
Hess Deep, Earth Planet. Sci. lett. 144 (1996) 123–132.
E. Hooft, H. Schouten, R.S. Detrick, Constraining crustal
emplacement processes from the variation in seismic layer
2A thickness at the East Pacific Rise, Earth Planet. Sci.
Lett. 142 (1996) 289–309.
C.R.B. Lister, Differential stress in the earth, Geophys. J.
R. Astron. Soc. 86 (1986) 319–330.
D. Curewitz, J.A. Karson, Geological consequences of dike
intrusion at mid-ocean ridges, in: Buck, W.R., Delaney,
P.T., Karson, J.A., Lagabrielle, Y. (Eds.), Faulting and Mag-
244
[49]
[50]
[51]
[52]
[53]
K.M. Gillis, M.D. Roberts / Earth and Planetary Science Letters 169 (1999) 227–244
matism at Mid-Ocean Ridges, Geophys. Monogr. Ser. 106,
American Geophysical Union, Washington, DC, 1998.
W.S.D. Wilcock, J.R. Delaney, Mid-ocean ridge sulfide deposits: Evidence for heat extraction from magma chambers
or cracking fronts?, Earth Planet. Sci. Lett. 145 (1996) 49–
64.
C.W. Burnham, Magmas and hydrothermal fluids, in:
Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore
Deposits, John Wiley, New York, NY, 1997, pp. 63–124.
B. Dutrow, D. Norton, Evolution of fluid pressure and fracture propagation during contact metamorphism, J. Metamorph. Geol. 13 (1995) 677–686.
R.P. Lowell, L.N. Germanovich, On the temporal evolution
of high-temperature hydrothermal systems at ocean ridge
crests, J. Geophys. Res. 99 (1994) 565–575.
C.J. Bruton, H.C. Helgeson, Calculation of the chemical
and thermodynamic consequences of differences between
fluid and geostatic pressure in hydrothermal systems, Am.
J. Sci. 283A (1983) 540–588.
[54] R.P. Dziak, C.G. Fox, A.E. Schreiner, The June–July 1993
seismo-acoustic event at CoAxial segment, Juan de Fuca
Ridge: evidence for a lateral dike injection, Geophys. Res.
Lett. 22 (1995) 135–138.
[55] P.T. Delaney, D.D. Pollard, Solidification of basaltic magma
during flow in a dike, Am. J. Sci. 282 (1982) 856–885.
[56] J.T. Wells, M.S. Ghiorso, Coupled fluid flow and reaction
in mid-ocean ridge hydrothermal systems: The behavior of
silica, Geochim. Cosmochim. Acta 55 (1991) 2467–2481.
[57] S.L. Brantley, The effect of fluid chemistry on quartz microcrack lifetimes, Earth Planet. Sci. Lett. 113 (1992) 145–
156.
[58] J.L. Bischoff, Densities of liquids and vapors in boiling
NaCl–H2 O solutions: a PVTX summary from 300º to
500ºC, Am. J. Sci. 291 (1991) 309–338.
[59] S. Sourirajan, G.C. Kennedy, The system H2 O–NaCl at
elevated temperatures and pressures, Am. J. Sci. 260 (1962)
115–141.