Capítulo 2: Marco Teórico 12
2.2. Desarrollo profesional, formación permanente y colaboración entre docente 19
2.2.3 La colaboración desde la formación, y en la práctica docente 23
The majority of mineralogical and petrographical data were obtained analysing
thirty-seven polished thin-sections from the Troodos Ophiolite CY-4 drill hole and thirty-
nine polished thin sections from the Upper Canada property. This was performed on a
Nikon Eclipse LV 100POL Microscope for the purpose of identifying primary and
secondary minerals. The microscope (Figure 3.2) is located in the Earth and Planetary
Materials Imaging and Analysis Laboratory at Western University.
The microscope has a 10X ocular lens and five objective lenses (5X, 10X, 20X,
50X and 100X) attached to the microscope. Photomicrographs were taken using a Nikon
Digital Sight DS-Ri1 high-resolution digital camera, located at the top of the microscope.
Nikon NIS- Elements software was used to store, edit and manipulate photographs and
the camera. Thin sections from the Troodos Ophiolite CY-4 drill hole were studied under
A B C
Figure 3.1. The sample preparation procedure includes: cutting, crushing, and milling using the equipment in the Rock Preparation Laboratory at Western University.
transmitted light with plane and cross polars, while representative thin sections from the
Upper Canada property were examined under both reflected and transmitted light. The
basic use of reflected light is for identification of opaque minerals, and their properties
(such as colour and colour intensity) influenced by reflectance.
3.3. X-Ray Diffraction (XRD) and Micro-X-Ray Diffraction (μ XRD)
X-ray diffraction (XRD) is a non-destructive technique for determination of mineral
phases typically in powdered sample. Micro X-ray diffraction (μ XRD) is a new versatile
Figure 3.2. Nikon Eclipse LV 100POL Microscope at the Earth and Planetary Materials Imaging and Analysis Laboratory, at Western University.
tool for the in situ examination of minerals, rocks, or polished thin sections on a
microscopic scale, allowing for comparison with other data or methods on a grain by
grain basis (Flemming, 2007).
This study used XRD patterns to define the alteration minerals in samples from the
CY-4 drill hole and the Upper Canada property. All selected whole rock powders (37
samples from the Troodos Ophiolite and 39 samples from the Upper Canada property)
were analysed using a Rigaku Rotaflex Powder X-Ray Diffractometer with
monochromated Co- Kα radiation (Co Kα, λ= 1.7902 Å) in the Laboratory for Stable
Isotope Science (LSIS) at Western University (Figure 3.3A).
This analytical procedure was used to estimate the quantity of each mineral as peak
intensity in an X-Ray diffraction pattern are proportional to the quantities of minerals that
compose a solid phase. The peak width in an X-Ray diffraction pattern is related to
precise determination of a crystal structure, size, and shape that compose the mineral.
These can provide valuable information about the analysed mineral (Dutrow and Clark,
2011). The analysed minerals were finely ground, homogenized, and an average bulk
mineralogy was determined.
The analytic procedure was based on the following steps: First, whole-rock powders
were homogenized by grinding to a fine powder (< ~10μm) with an agate mortar.
Powders were mounted on a sample holder and detected by a rotating detector. X-ray
diffraction data were collected at 160 kV and 45 mA, from 2° to 82° 2Ɵ with a step size
of 0.02°, at a speed of 10° per minute when the powder is struck by the X-ray beam.
were characterised using the Bruker AXS EVA software package (BrukerAXS, 2006),
and correlated with the International Center for Diffraction Data (ICDD) PDF4 database.
The final results of integrated diffractograms with ICDD pattern models for both
investigated localities are documented in Appendices 1 and 2.
A total of five selected samples (four diabase dyke samples and one sample of
gabbro from the Troodos Ophiolite Complex) were selected for micro X-ray diffraction
(μXRD) analysis. μXRD was performed on the Bruker –AXS D8 Discover micro X-ray Diffractometer in the X-Ray Diffraction and Microdiffraction Laboratory at Western
University (Figure 3.3B), to correlate with x-ray diffraction patterns from the Rigaku
Rotaflex Diffractometer. The Bruker–AXS D8 Discover Diffractometer uses a sealed
A B
Figure 3.3: The X-Ray Diffraction machines: A- The Rigaku Rotaflex Powder X- Ray Diffractometer with monochromated Co- Kα radiation. Major parts of the diffractometer are: an X-ray tube, a sample holder, and an X-ray detector at the LSIS; and B- The Bruker –AXS D8 Discover micro X-ray Diffractometer in the X-Ray Diffraction and Microdiffraction Laboratory. Parts of the Diffractometer are : a) scaled Cu source, b) Gobel mirror with parallel beam optics, c) snout collimator, d) optical monitor with camera, e) laser, f) general area detector diffraction system, g) stage, h) sample mount with garnets, and i) sample as viewed under the microscope (from Flemming, 2007).
tube source with accelerating voltage of 40 kV and 40 mA, which produces Cu Kα
radiation (Cu Kα, λ=1.5418 Å). A gobel mirror with parallel beam optics was used in order to remove Kᵦ radiation and maximize diffraction beam intensity for non-flat samples. An exchangeable pinhole collimator snout produced a beam diameter of 500 μm
or 50 μm. The source and the detector were mounted on a goniometer (Ɵ - Ɵ geometry which allows the sample to remain stationary), while diffraction rays were detected by a
general area diffraction system (GADDS) with 15.2 cm in diameter (Flemming, 2007).
Each crystal diffracts a unique set of 2Ɵ angles, which creates the characteristic Debye
rings. Sample preparation is the same as for the Rigaku Rotaflex Diffractometer. Whole
rock powders were then placed in glass sample holders on the XYZ stage for analysing.
The samples were precisely targeted using an optical microscope monitor and digital
video camera system (Debye Scherrer camera with digital 2-D detectors). The laser,
microscope optics and X-Ray beam were aligned when the laser was in the optical focus
with the microscope crosshairs, meaning that the area of interest on the sample was at the
center of diffraction. The angles between the source and the sample (Ɵ1) and the angle
between the sample and detector (Ɵ2) were measured, allowing the angle between the
detector and the source to move independently (by the operator) about the sample stage.
Also, changing the angle between them means the Bragg’s Law conditions are satisfied:
Ɵ1+ Ɵ2 = 2Ɵ and this relationship is always maintained (Flemming, 2007). For analyzing samples from CY-4, off-coupled scanning was chosen as an appropriate model with no
oscillation on samples, as suggested by Dr. Roberta Flemming. A few parameters were
defined: Ɵ1=6, Ɵ2 =16.5, width =28.5, and analysis time of 60 minutes per frame. The
in order to identify the mineral phases from the International Center for Diffraction Data
PDF4 database.
3.4 Introduction to Oxygen Stable Isotopes
Atoms of the same element with the same number of protons and electrons, and a
different number of neutrons are called isotopes. Naturally occurring oxygen is composed
of three stable isotopes: 16O (99.758%), 17O (0.039%) and 18O (0.205%) (Hoefs, 2004). It
occurs in gaseous, liquid, and solid states. Because of higher abundances of 18O and 16O,
we used the ratio of 18O to 16O to determine the oxygen isotopic composition of a rock.
This is expressed as:
δ 18
O= [(18O/ 16 O) sample / (18O/ 16 O) VSMOW -1] x 1000,
where (18O/ 16 O) sample represents the measured ratio of heavy to light oxygen isotopes in
a rock sample (Clark and Fritz, 1997). The oxygen isotopic data for the rock sample (δ
18
O) is reported relative to VSMOW (Vienna Standard Mean Ocean Water), in the
standard ‘per mill’ notation (‰). The oxygen isotope composition of the rock sample depends on the 18O-contents of the constituent minerals and the mineral proportions. The
18
O-contents of the constituent minerals depend on bond-type and strength in the crystal
structure (Hoefs, 2004).
Equilibrium and kinetic effects can cause isotope ratios of oxygen to change due to
chemical and physical properties between the isotope, and the oxygen. This phenomenon
chemical equilibrium is called the isotope fractionation factor. The isotope fractionation
factor is defined as:
α(A-B) = RA/RB
where RA is the ratio of heavy to light isotope in phase A, and RB is the ratio of heavy to
light isotope in phase B (Faure, 1986).
3.5 Previous Work on Oxygen Stable Isotopes in Hydrothermal Systems
The stable oxygen isotope composition of a rock has become an important tool in
modern petrology. Fractionation of oxygen isotope ratios is useful to provide information
on different problems such as: 1) the conditions and mechanisms of minerals and rock
formation; 2) the origin and evolution of magmas; 3) the interaction between magma and
country rocks; 4) the nature of the fluids involved in geological processes; 5) the nature of
the fluids and possible fluid migration during metamorphism; and 6) interaction between
rocks and circulating fluids (Turi, 1988).
The temperature dependence of oxygen isotopic fractionation factors between
minerals in rocks and fluids provides a crucial key in understanding hydrothermal
alteration processes within the oceanic crust. Oxygen is the principal element in most
rocks. Interaction between rocks and circulating fluids can result in recrystallization of
primary minerals in rocks or the precipitation of new minerals. This process depends on
the 18O content of the original minerals (strength, and bond type in their crystal structure), and the proportions of these minerals in the rocks as well as any interaction
have focused on the importance of rock-fluid interactions from aspects of the kinetics and
oxygen isotope exchange mechanisms. There are three possible mechanisms that can
cause a shift in the oxygen isotope ratio between existing minerals and fluids: 1) solution-
precipitation (when large crystals grow at the expense of smaller ones); 2) chemical
reaction (when the break-down of original crystals occurs with the formation of new
ones); and 3) diffusion (when isotopic exchange happens between the crystals and the
fluids with no change in morphology of original crystals); (O’ Neil and Taylor, 1967; Taylor, 1974; Matthews et al., 1983b, 1983c; Giletti, 1985; Criss et al., 1987; Gregory et
al., 1989).
In order to determine rock-seawater exchange, many authors use mid-ocean ridge
basalt (MORB) 18O values as indicators of the oxygen isotope composition of the mantle (Muehlenbachs and Clayton, 1972a; Muehlenbachs and Clayton, 1976; Muehlenbachs
Figure 3.4. The oxygen stable isotope variations of important geological reservoirs in nature (from Hoefs, 2004).
1986; Mattey et al., 1994, Harmon and Hoefs, 1995). MORBs have a homogeneous
reservoir for oxygen with a mean δ 18O value of 5.7± 0.2 ‰ (Clayton and Muehlenbachs,
1976; Muehleubachs, 1998). During partial melting of the mantle, isotope fractionation is
small because of the high temperatures and the oxygen isotope composition of MORB
remains similar to the mantle source. However, weathering and extensive oxygen isotopic
exchange between the oceanic crust and seawater fluids complicates the situation (White,
2005). The δ18O value of seawater is buffered by weathering or hydrothermal alteration
processes (at temperatures from 250 to 350°C) at mid-ocean ridges to the present δ 18O
value of 0± 2 ‰, with no tendency for the isotopic composition of seawater to change
(Clayton and Muehlenbachs, 1976; Muehleubachs, 1998). This means that the oxygen isotope composition of MORB on the seafloor are not in equilibrium with seawater and
will therefore interact at either high or low-temperature. The changes will result in a shift
in δ18
O values of MORBs as they become depleted or enriched relative to the primary
oxygen isotope value of MORBs. In addition, low-temperature alteration results in the
development of secondary phyllosilicate minerals leading to 18O-enrichment (in clay,
carbonate, serpentine, zeolite and chlorite groups; Clayton and Muehlenbachs, 1972). In
contrast, high-temperature alteration includes the development of secondary minerals like
epidote groups leading to 18O-depletion relative to unaltered basalts (Clayton and
Muehlenbachs, 1972). Oxygen isotopic values from ophiolites and greenstone belts are
important tracers of this exchange process between rocks and hydrothermal fluids
(Gregory, 2003). Stable isotope studies on ancient ophiolites in different tectonic settings
provide direct evidence of seawater and rock interactions (Muechlenbachs, 1998, Alt and