1811-5209/11/0007-0327$2.50 DOI: 10.2113/glements.7.5.327
INTRODUCTION
Tourmaline has precily the characteristics of a forensic mineral: it is widespread, occurring in most rock types; it is stable over a significant portion of the pressure (P)–temperature (T)–composition (X) space of the crust (Dutrow and Henry 2011 this issue); its crystal structure accommodates an exceptional range of geochemically different elements (Hawthorne and Dirlam 2011 this issue); its (trace) element and isotopic compositions are responsive and refl ective of physical and chemical conditions encoun-tered and of element source-rervoirs (e.g. Marschall and Jiang 2011 this issue; Slack and Trumbull 2011 this issue); and it shows negligible intracrystalline element diffusion over geologic time, even at temperatures in excess of 600 ºC (van Hinsberg et al. 2011). Becau of the c
haracteristics, tourmaline records a signature of the conditions and chem-istry of its environment, in an extraordinary range of ttings within the Earth, and, most crucially, prerves this signature for later interrogation. Moreover, tourmaline grains are commonly growth-zoned, and this “memory” is recorded and stored for the full growth history of the grain, just as tree rings store information about the tree’s environment during its growth. Unconformities can develop in this record, when growth stalls for lack of boron (B) or growth zones are eroded away, but the can be informative as well (e.g. van Hinsberg and Schumacher 2011).
The primary microstructures and compositions of tourma-line are commonly well prerved, even when grains are weathered from their parent rocks and transferred to di-
ment. When the are properly
interpreted, they provide robust
evidence that tracks the mineral’s
formation and modifi cation by
igneous, dimentary, and meta-
morphic process. In this article,
we explore a hypothetical record
of a single, multigenerational tour-
maline grain and illustrate how
tourmaline can record geologic
information in igneous, dimen-
tary, and metamorphic ttings
(F IGS. 1, 2). This “tourmaline
construct” provides a framework
for discussion of tourmaline
formation and modification in
the varied geologic ttings.
TOURMALINE IN IGNEOUS ROCKS
Many tourmalines begin “life” in the geologic record as crystallization products of an igneous melt. The most common tourmaline-bearing igneous rocks are tho that have undergone signifi cant chemical fractionation, becau boron, an esntial constituent of tourmaline, behaves incompatibly in igneous systems and is therefore highly enriched in fractionated melts (e.g. London et al. 1996). Tourmaline is most important in silicic and peraluminous volcanic and plutonic rocks, such as rhyolites, granites, granitic pegmatites, and nodular granites. If the rocks are derived from the melting of metadimentary material, which is commonly enriched in B compared to average crust, tourmaline may occur in substantial modal amounts. Tourmaline displays a myriad of morphologies in igneous deposits, from beautiful faceted prisms (e.g. Pezzotta and Laurs 2011 this issue), to skeletal grains, to interstitial mass, and to mineral fibers. The shapes reveal the timing and style of tourmaline growth in the melt. Commonly, tourmaline forms late in the crystallization history of the magma, and its growth may be rapid owing to the enhanced concentration of boron and water in the residual melt.
Chemically, igneous tourmaline tracks the composition and mineral asmblage of its host magma. Therefore, it provides a fi rst-order approximation of the composition of the magma from which it nucleated, including changes in major components such as Fe, Mg, and Mn (e.g. Jolliff et al. 1986). Tourmaline in intermediate igneous rocks gener-ally has compositions that are Na-rich and intermediate in Fe and Mg contents, e.g. schorl and dravite, whereas tour-maline crystallized in moderately fractionated granitic rocks tends to be Fe-rich (i.e. schorl and foitite; e nomen-clature in Hawthorne and Dirlam 2011). With incread fractionation, tourmaline gets enriched in Li, so that in
T ourmaline is nature’s perfect forensic mineral. From a single grain, the full geological past of its host rock can be reconstructed, including the pressure–temperature path it has taken through the Earth and the changing fl uid compositions it has encountered. Tourmaline is able to provide this record owing to its compositional and textural nsitivity to the environ-ment in which it grows, and is able to prerve this record becau element diffusion in its structure is negligible. Furthermore, tourmaline has an excep-tionally broad stability range, allowing it to record conditions in igneous, dimentary, metamorphic, and hydrothermal ttings. As our mineralogical and geochemical tools advance, we are able to interrogate tourmaline’s memory with increasing precision, making tourmaline a truly powerful indi-cator of conditions in the Earth.
K EYWORDS: tourmaline, petrogenetic indicator, compositional zoning,
growth history
Vincent J. van Hinsberg1, Darrell J. Henry2, and Barbara L. Dutrow2
1 Department of Earth Sciences, University of Oxford
South Parks Road, Oxford, UK
E-mail: V.J.
2 Department of Geology and Geophysics, Louisiana State
University, Baton Rouge, LA 70803, USA Photomicrograph of an hourglass ctor zoned
tourmaline in a metapelite (blue +c ctor, brown –c ctor and yellow a ctor)
highly differentiated pegmatites, the Li species elbaite, fluor-liddicoatite, and rossmanite are characteristic. Tourmaline is most common in the B-enriched Li–Cs–Ta family of pegmatites and occurs more rarely in Nb–Y–F pegmatites (e.g. London 2008). Volatile-rich, commonly late-stage cav
ities in pegmatites can contain gem varieties (e Pezzotta and Laurs 2011).
it的宾格
Becau the environment continuously changes in a differ-entiating magmatic system, tourmaline commonly exhibits well-developed color and compositional growth zoning, in addition to ctor zoning (F IG. 1B STAGE 1; F IGS. 2, 3; Hawthorne and Dirlam 2011). Chemical trancts across such zoned grains provide a record not only of the evolving melt composition but also of dissolution/precipitation events and changes in alkalinity, fluorine content, and redox state of the melt and associated fl uid (e.g. Dutrow and Henry 2000; London 2008).
医患关系的现状
The late-stage hydrothermal fl uids associated with many igneous plutonic rocks are enriched in B and can infi ltrate the surrounding host rocks. This infl ux of boron may cau tourmaline to form, creating tourmaline-rich rocks at the margins of igneous bodies and in their country rocks. Such fl uids proximal to igneous bodies can be highly oxidized, yielding tourmaline with a signifi cant proportion of Fe3+ (e.g. povondraite), and they are associated with many types of ore deposits (Slack and Trumbull 2011). Metasomatic tourmaline generally exhibits fi ne-scaled oscillatory zoning and typically displays a hybrid composition that refl ects reaction between the host rock and a B-bearing fl uid pha. One can commonly determine the relative contributions of the host rock and fl uid to the metasomatic tourmalines from their isotopic and trace element compositions, as well as
reconstruct changes in the contributions over time. Tourmaline thus tracks the evolution of magmatic–hydro-thermal systems, and it will retain this information despite paration from its site of crystallization by weathering. TOURMALINE IN SEDIMENTS
AND SEDIMENTARY ROCKS
Tourmaline as a Detrital Mineral
and Provenance Indicator
The exceptional mechanical and chemical stability of tour-maline makes it highly resistant during weathering. Once expod to surface conditions, tourmaline can be disag-gregated from its igneous, metamorphic, or dimentary parent rock to become a clastic component in diment. Together with zircon and rutile, tourmaline is one of the most durable heavy minerals in dimentary environments.
Tourmaline in diment retains the signature of its original host rock (F IG. 1A, 1B STAGE 2) and has, therefore, long been a provenance indicator in clastic dimentary rocks. Early attempts focud on features such as absorption color, grain size, and degree of roundness either to infer provenance or to correlate clastic dimentary units (e.g. Krynine 1946). However, absorption colors
can yield equivocal results in source-rock correlations. The compositions of detrital tourmaline grains are much better indicators of likely source rocks. “Environmental diagrams” that compare Al–Fe–Mg and Ca–Fe–Mg abundances show how detrital tourmaline compositions relate to probable source rock types (F IG. 4; Henry and Guidotti 1985). The diagrams, along with other compositional and textural variables, successfully yield provenance determinations in a number of different lithologic ttings (e.g. Henry and Dutrow 1992; Morton et al. 2005). In addition, the relative degrees of rounding of tourmaline clasts provide general information on the energetics and potential cycles of di-mentation that each clast has undergone (Krynine 1946).
A challenge in using tourmaline composition to derive diment provenance is that composition is not solely refl ective of host rock. Variations in pressure, temperature, and oxygen fugacity, among others, can modify the parent-rock signature. The prence of compositional zoning in
stage 1
magmatic
stage 2
dimentary
stage 3
diagenetic
stage 4
metamorphic
stage 5
metasomatic
stage 6
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hydrothermal
F IGURE 1(A) Schematic drawing of an “idealized” multigenera-
tional tourmaline that encountered various geologic process and environments (SEE F IGURE 1B). From the inside outward, the grain contains: a rounded detrital core that formed by crystallization from a melt; asymmetric diagenetic overgrowths with preferred growth towards the +c pole; metamorphic and metaso-matic zones overgrowing an internal foliation (s1) marked by inclu-sions; and a fracture aled by hydrothermal growth on the fracture surfaces. Polar growth and hourglass ctor zoning are marked by color shading. The heavy black line marks the tranct described in F IGURE 2. (B) Geologic ttings for growth of the tourmaline grain shown in A. The are: magmatic tting (stage 1) with subquent weathering (stage 2); diagenetic growth on relict grains in a di-mentary basin (stage 3); prograde metamorphic growth (stage 4); retrograde metasomatic growth owing to input of boron in fl uids from a cooling pluton (stage 5); and growth in extensional brittle fractures associated with ore formation (stage 6).
B
detrital grains provides a further challenge. However, such zoning also provides an opportunity, beca
u it allows trends in composition to be matched to a source region. Trends are more distinctive than a single composition, thereby improving source identifi cation, and may even be ud to pinpoint a specifi c location within a source terrane. For example, if the model tourmaline in F IGURE 1 were to end up as a detrital grain, the P –T and compositional history retrieved from it would provide a highly charac-teristic signature for u in provenance studies.
Diagenetic Tourmaline
The burial and lithifi cation of dimentary detritus to form dimentary rock commonly lead to new tourmaline growth on detrital grains or nucleation of a new tourmaline generation. A diagenetic origin is inferred for the occur-rences becau overgrowths typically develop as fragile, slender, syntactic needles (F IG . 5) that would not have survived transport in a clastic dimentary environment (Krynine 1946). Most commonly, such tourmaline occurs as monopolar (restricted to one side of the grain) or highly asymmetric overgrowths on detrital tourmaline (e.g. Krynine 1946; Sperlich et al. 1996; Henry and Dutrow 1996), with distinctly different colors for overgrowths on either side of their detrital core (commonly blue and brown). A pale-hued overgrowth is found on the +c end of the detrital grain, whereas a darker-colored overgrowth develops on the –c end (e Hawthorne and Dirlam 2011). Growth also predominates on, and may even be restricted to, the +c s
关于母爱的段落ide, which, together with color, provides a good marker of the crystallographic orientation of rounded detrital grains (F IGS . 1B STAGE 3; F IG . 5). The distinctive over-growths on the two sides of a detrital grain testify to tour-maline’s polar nature (e Hawthorne and Dirlam 2011), which results in tourmaline growing in the +c direction taking up different elements than tourmaline growing in the –c direction. This feature is uful, becau the polar differences are temperature dependent and can be ud to constrain the temperature of diagenetic tourmaline forma-tion (e.g. Henry and Dutrow 1996).
In contrast, diagenetic tourmaline with no evidence of a detrital tourmaline nucleus is common in carbonate rocks, stromatolites, and, rarely, in mudstones (Srivastava and Schnitzer 1976). Compositions of pale-colored, diagenetic
tourmaline are generally X -site-vacant foitite or magnesio-foitite species (e.g. Henry et al. 1994; Ronberg and Foit 2006). The diagenetic tourmaline grains continue to grow as temperature and pressure increa and B is relead from associated minerals.
TOURMALINE IN METAMORPHISM
Metamorphic rocks, especially tho of metapelitic compo-sition, are the cond most important host
of tourmaline. In contrast to their igneous counterparts, many metamor-phic tourmaline grains develop on preexisting detrital grains or diagenetic cores, as is shown in the model tour-maline (F IG . 1A ). Tourmaline textures and compositions refl ect both local metamorphic reactions and changing P –T –X conditions, and this chemical and textural evidence of events can be prerved from the lowest to the highest grades of metamorphism (F IG . 1A , 1B STAGES 4 AND 5; F IG . 2). During progressive metamorphism, tourmaline grows when B is available and the composition of the fl uid is amenable to tourmaline formation (e.g. Dutrow et al. 1999). In diments and metadimentary rocks prior to the formation of tourmaline, B predominantly resides in layer silicates, with concentrations up to veral hundred parts per million (Grew 1996; Steppan 2003). Even after tourmaline begins to grow, coexisting phas retain appre-ciable amounts of B. In a typical pelitic metadiment (Steppan 2003), the B distribution among coexisting phas is in the order tourmaline >> muscovite > andalusite > chlorite > biotite > plagiocla > garnet > kyanite > stau-rolite, although it varies as a function of metamorphic grade. In general, the sheet silicates contain more B than their breakdown products, leading to relea of B during prograde metamorphism.
The relea of B along a typical prograde metamorphic path (e.g. F IG . 2) can be predicted by combining thermo-dynamic calculations of changes in mineral proportions for a metapelitic bulk com
position (using the PerpleX suite of programs; Connolly 2005) with typical B contents for metamorphic minerals (e.g. Steppan 2003). The calcula-tions suggest a steady, but punctuated, supply of B on the prograde metamorphic path, with the largest B relea at low grade by breakdown of muscovite and chlorite (F IG . 6). Conquently, tourmaline develops episodically as B is
s t g t + b t +
s i l
T emperature (°C)
P r e s s u r e (k b a r )
c a l c i t e -o u
t
s t + b t 0
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i t e
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p y r o p h y l l i t e
k y a n i t e
g t + c h l + m s
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Ca (aq)
Na (aq)
T emperature (°C)
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C o n c e n t r a t i o n i n fl u i d
Distance
C o n c e n t r a t i o n i n t o u r m a l i n e
Na Ca
Zn growth in fracture
Li core rim
calcite out staurolite out
staurolite in growth in fracture
core
rim
2
P –T conditions (A ) and element
contents (B , C ) for lected major and trace elements along a core–rim tranct of the tourmaline grain shown in F IGURE 1A , as well as the fracture growth (shown as a star in A). Arrows show the time direction. The spikes in composi-tion (points 1–3, F IG . 2B ) allow for mineral reactions to be recog-nized and for a temperature to be assigned to each reaction using inter-ctor tourmaline thermom-etry. Combined with reaction positions from the literature (the green and blue lines in the P –T diagram), this allows the P –T path to be constrained (in red; stip-pled parts of the P –T path were not recorded for lack of tourma-line growth). Tourmaline compo-sitions, combined with partition coeffi cients from the literature, allow Ca and Na contents in the fl uid to be reconstructed, shown sc
hematically against temperature along the P –T path in F IGURE 2C .
relead in mineral reactions, and the timing and amount of tourmaline growth are linked to changes in the asm-blage of associated minerals. A rock dominated by musco-vite will relea its B load at different times along the metamorphic path compared to one dominated by chlorite, becau of differences in their respective stability. Boron relead in a mica-rich unit, which releas aqueous fl uid upon mica decomposition, is likely to infi ltrate enclosing rocks, where it may form distinct tourmaline overgrowths or lead to mineral replacement and/or development of tourmaline-rich pudomorphs (Dutrow et al. 1999). The liberation of B is nearly complete at peak metamorphic conditions such that, for a clod system, tourmaline is unlikely to grow on the retrograde path (F IG . 6). Comparison of the thermodynamically calculated B relea with the timing of growth of tourmaline from three different metapelites from the Massif Central, France (colored lines in F IGURE 6) shows good overall agreement and matches the prediction that the most substantial growth takes place at lowest grade.
The composition of metamorphic tourmaline changes notably throughout its growth history as it directly refl ects the changing environment (F IG . 2). Nonetheless, typical compositions are primarily schorl to dravite (i.e. the Fe 2+ to Mg ries), albeit with appreciable foitite, olenite, and uvite compo
nents, which introduce X -site vacancies, Al on the Y site and Ca on the X site, respectively. With increasing grade, the Mg/Fe 2+ ratio and the content of Al on the Y site increa, along with vacancies and Ca content on the X site (van Hinsberg et al. 2011). However, the general trends are easily disturbed. For example, where sulfi des quester most of the Fe, anomalously Mg-rich tourmaline ensues (Henry and Dutrow 1992). Whereas the Mg/Fe 2+ ratio of tourmaline commonly shows a smooth change from core to rim with changing conditions, other elements, including Ca, show spikes and troughs that refl ect sudden changes in element abundance as the mineral asmblage or fl uid composition changes (e.g. calcite breakdown; F IG . 2B , C ).
Constraining Conditions of Tourmaline Formation
Recently, it has become possible to directly constrain the temperature of tourmaline formation on the metamorphic path by using tourmaline geothermometry. The most promising approach is the temperature-dependent parti-tioning of elements among ctors in a single tourmaline grain that displays hourglass-shaped ctor zoning, i.e. different compositions in a single growth horizon for domains that originated on different growth surfaces (van
Hinsberg and Schumacher 2007). This inter-ctor ther-mometry can be applied to isolated grains, including
detrital clasts and tourmaline inclusions, and to single growth zones. Its greatest value lies in application to tour-maline grains that show both ctor and growth zoning, becau in the cas complete temperature histories can be extracted from single grains (e.g. van Hinsberg and Schumacher 2007, 2011). Moreover, tourmaline thermom-etry access the low-temperature prograde conditions of metamorphic growth, for which few other robust mineral thermometers are available.
In ctions cut parallel to tourmaline’s long axis, hourglass zoning occurs as a pale-colored, commonly blue ctor in the +c direction, a dark-colored, typically dark brown ctor in the –c direction, and an intermediate-colored ctor on either side (F IG . 1A ). In ctions perpendicular to the long axis, tourmaline exhibits a homogeneous, light- or dark-colored core surrounded by a zoned rim. This core –rim color difference can be mistaken for two-stage growth. Sector zoning refl ects tourmaline’s polar nature and is a different manifestation of polar diagenetic growth. Although tourmaline ctor zoning is commonly regarded as a diquilibrium feature, this is a misconception. The ctor differences form as a result of different local equi-libria between the various growth surfaces and their iden-tical host environment (van Hinsberg and Schumacher 2007). There is no direct exchange of elements between the ctors; instead, they communicate through the fl uid, an exchange that equals direct exchange in thermody-namic terms.
Sector zoning and polar growth are controlled by differ-ences in the surface properties of the growth surfaces; differences in charge are of most importance in tourmaline. Becau the Si-tetrahedra all point in the same direction along the c axis in the tourmaline structure (Hawthorne and Dirlam 2011), a partial positive charge develops on surfaces at high angles to c . A corresponding negative charge develops on the opposite side, leading to the pref-erential incorporation of higher-charged cations for growth in this direction, such as Ti 4+ on the Y site and Ca 2+ on the X site. With increasing temperature, the ability of time-equivalent, but crystallographically distinct, growth
surfaces to gregate elements decreas becau the incread vibration of the atoms diminishes the charge asymmetry along the c axis; the inter-ctor partition coef-fi
cients converge to unity. This change in partition coef-fi
cients provides the basis for thermometry. Becau the compositional contrasts are in equilibrium only at the
growth surface, ctor zoning should disappear by element F IGURE 3
Polished slab of color-zoned tourmaline cut perpen-dicular to the c axis from a pegmatite in Madagascar.
The tourmaline exhibits color zoning related to both ctor zoning
(cross in core) and oscillatory zoning (rim).
F IGURE 4Environmental diagram correlating tourmaline compo-sition to host-rock type. The locations of lected
tourmaline species are indicated. The diagram is modifi ed after
Henry and Guidotti (1985).
diffusion in the crystal bulk. The prervation of ctor zoning is therefore not only evidence of slow diffusion in tourmaline, but also a confi rmation that a grain has retained its chemical signature.
At prent, other tourmaline thermometers only provide
broad indications of conditions. The scatter obrved in
calibration attempts ems to refl ect disorder of cations
over the different tourmaline sites, in particular Fe, Mg,
and Al over the Y and Z sites (Hawthorne and Dirlam 2011),
and changes therein with changing conditions (van
Hinsberg and Schumacher 2009). Element exchange and
occupancy on the X site is more promising, especially
Na –Ca exchange between tourmaline and plagiocla (van
Hinsberg and Schumacher 2009) and X -site occupancy
(Henry and Dutrow 1996; Ertl et al. 2010).
Constraining the pressure of formation is more diffi cult.
Tourmaline barometers have not as yet been established, although tourmaline’s K/Na ratio increas with pressure in exchange with biotite or phengite (van Hinsberg, unpub-lished data). Nonetheless, pressures can be constrained indirectly, for example, by barometry on mineral or fl uid inclusions. In addition, specifi c mineral reactions are en as sharp changes in tourmaline composition. If the position of such a reaction is known in P–T space and the tempera-ture can be constrained from tourmaline thermometry, pressure can be established from the interction of this reaction line in P–T space (e.g. van Hinsberg and Schumacher 2011). This approach is highlighted for our
model tourmaline, in which the terminal breakdown of carbonate appears as a spike in Ca (F IG . 2B ), whereas the reaction garnet + chlorite + muscovite to staurolite + biotite + quartz + fl uid is en by the combination of a trough in
Zn + Li and a spike in Mn + Ca owing to the different element preferences of products and reactants. By combining this information, the P–T path of a tourmaline-bearing rock can be established (F IG . 2A ).
Our model tourmaline shows both growth and ctor zoning (F IG . 1A ), allowing reconstruction of its thermal history. Compositions within the same growth zone, but on either side of the ctor boundary, provide an inter-ctor partition coeffi cient that constrains the formation temperature of this growth zone. The calculations can be performed along the entire length of the crystal. From core to rim, the compositional differences between the ctors decrea, which shows that this quence corre-sponds to prograde growth (growth with increasing temperature). Qualitatively, this decrea in compositional contrasts among ctors is en as a decrea in color contrast, and the peak growth zone is identical in color for all ctors. In natural metamorphic tourmalines, this convergence is obrved for temperatures in excess of 650 °C (van Hinsberg and Schumacher 2007).
Growth of tourmaline during cooling and exhumation (retrograde growth) is unlikely in a clod system, becau B has already been questered in tourmaline or lost to the fl uid upon reaching peak metamorphic conditions (F IG . 6).
Metasomatic introduction of B is the most common cau of tourmaline growth on the retrograde path, although such metasomatism can also take place during other stages of the rock’s formation. The typical retrograde rehydration obrved in metamorphic rocks and the high solubility of
B in hydrothermal fl uids suggest that this process is
common. Intrusion of late “suturing” granites in orogenic
belts and the likelihood of relea of fl uids from such
plutons is a common source for retrograde boron addition,
as is shown for our model tourmaline (F IG 1B STAGE 5). In
many cas, this retrograde overprint takes place over a
range of temperature, and tourmaline may record this
history (van Hinsberg and Schumacher 2007). The growth
of metamorphic tourmaline can thus span a remarkable
range of conditions and provide a treasure trove of infor-mation on the P–T history of its host rock.
Tourmaline as an Indicator of Fluid Composition
Tourmaline not only prerves P–T information but also captures a fi ngerprint of the composition of its growth environment(s). This fi ngerprint allows changes in mineral paragenesis to be recognized, but its real power is the record provided on trace element mobility and the isotopic signa-ture of the host environment. For example, in h ydrothermal
ttings this fi ngerprint of element mobility has allowed tourmaline to be ud as an exploration tool in the arch for ore deposits (e.g. Slack and Trumbull 2011). Recent experimental work on element partitioning between
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tourmaline and fl
uid permits a quantitative reconstruction of element contents in the coexisting aqueous fl
uid (von Goerne et al. 2011 and references therein). In a way that F IGURE 5
Detrital tourmaline grain with asym-metric diagenetic tourmaline over-growth. (A ) Optical photomicro
graph highlighting the color distinction between the detrital core (blue) and overgrowth (yellow; top is +c ).
(B ) Backscattered-electron image illustrating the textural and chemical complexity of the diagenetic overgrowth shown in A.
A B
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T emperature (°C)
蚂蚁外形特点B r e l e a s e d (%)
100t o u r g r o w t h (%)
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predicted B relea
% growth of 3 natural
tourmaline grains
F IGURE 6
For a typical prograde path (F IG . 2A ), B is liberated as grade increas (black line), as a result of decreasing
modal abundance of the sheet silicates. This progressive, but episodic, B relea agrees well with the growth history of three natural metamorphic tourmaline grains (colored lines; data from van Hinsberg and Schumacher 2011).
