GeolCarp_Vol67_No5_417

GEOLOGICA CARPATHICA

, OCTOBER 2016, 67, 5, 417 – 432

doi: 10.1515/geoca-2016-0026

www.geologicacarpathica.com

Age and origin of fluorapatite-rich dyke from Baranec Mt.

(Tatra Mts., Western Carpathians): a key to understanding

of the post-orogenic processes and element mobility

ALEKSANDRA GAWĘDA



, KRZYSZTOF SZOPA

, DAVID CHEW

, URS KLÖTZLI

AXEL MÜLLER

, MAGDALENA SIKORSKA

and PAULINA PYKA

Faculty of Earth Sciences, University of Silesia ul. Będzińska 60, 41-200 Sosnowiec, Poland;



aleksandra.gaweda@us.edu.pl

Department of Lithospheric Research, University of Vienna, Althanstrasse 14, 1090 Vienna, Austria

Department of Geology, School of Natural Sciences, Trinity College Dublin, Dublin 2, Ireland

Natural History Museum, University of Oslo, P.O. Box 1172, Blindern, 0318 Oslo, Norway

Polish Geological Institute-Polish Research Institute, ul. Rakowiecka 4, Warsaw, Poland

(Manuscript received January 10, 2016; accepted in revised form September 22, 2016)

Abstract: On the southeastern slope of the Baranec Mount in the Western Tatra Mountains (Slovakia) an apatite-rich

pegmatite-like segregation was found in the subvertical fault zone cutting metapelitic rocks. Two zones: felsic (F) and

mafic (M) were found, differing in mineral assemblages and consequently in chemistry. Fluorapatite crystals yield

a LA-ICP-MS U-Pb age of 328.6 ± 2.4 Ma. A temperature decrease from 634 °C to 454 °C at a pressure around 500 to

400 MPa with oxygen fugacity increasing during crystallization are the possible conditions for formation of the

pegmatite-like segregation, while secondary alterations took place in the temperature range of 340 – 320 °C. The Sr-Nd

isotope composition of both apatite and whole rock point toward a crustal origin of the dike in question, suggesting

partial melting of (P, F, H

O)-rich metasedimentary rocks during prolonged decompression of the Tatra Massif.

The original partial melt (felsic component) was mixed with an external (F, H

O)-rich fluid, carrying Fe and Mg fluxed

from more mafic metapelites and crystallizing as biotite and epidote in the mafic component of the dyke.

Key words: dyke, apatite, U-Pb apatite age, Tatra Mountains.

Introduction

Syntectonic veins cutting metamorphic complexes are

important in deciphering the geological processes acting

during uplift and decompression in collisional orogens

(Druguet et al. 2008; Chen et al. 2012). Such veins, being

a result of channelized fluid/melt flow, are discordant or

parallel to host rock foliation and could be classified as peg-

matites (Liebscher et al. 2007; Chen et al. 2012). In particu-

lar, the mineral assemblages and geochemical features of

pegmatite-like mineral segregations carry information about

the partial melting processes, the origin and evolution of the

melts and fluids during plate subduction, new melt genera-

tion and the interaction between fluid, melt and country rocks

during decompression (e.g., Miller et al. 2002; Schmidt &

Poli 2003; Lü et al. 2012). The mineral composition of such

veins is simple (feldspar, quartz, micas, amphiboles,

pyroxenes, epidote-group minerals in different proportions,

with accessories like garnet, zircon and apatite; Chen et al.

2012; Lü et al. 2012) and they do not show the typical peg-

matite zonation (Simmons & Webber 2008).

Fluorapatite-rich veins are rarely found in the Tatra Moun-

tains (Fig. 1a,b) and were subdivided as a special type, dif-

fering from the typical muscovite-type pegmatites (Gawęda

1993, 1995). The muscovite-type pegmatites were connected

to a muscovite dehydration melting process, and dated by

whole-rock Rb-Sr method at ca. 345 Ma (Gawęda 1995). On

the southeastern slope of the Baranec Mount (Western Tatra

Mountains; Fig. 1b,c) an apatite-rich pegmatite-like dyke

was found in a subvertical fault zone, which also hosts an

eclogite boudin and is bordered by strongly mylonitized

metapelitic rocks (Fig. 1c). The aim of this paper is to deci-

pher the origin and age of the apatite-rich pegmatite-like

dyke. Apatite LA-ICP-MS U-Pb dating was used to constrain

the timing of crystallization and decompression activity

which is then compared to published data from the crystal-

line basement of the Tatra Mountains.

Geological setting and sample description

The Tatra Mountains represent the northernmost crystal-

line core of several “core mountains”, present within the

realm of the Central Western Carpathians (Fig. 1a,b). The

crystalline core of the Tatra Mountains comprises a Variscan

polygenetic granitoid pluton and its metamorphic envelope

(Fig. 1b; Morozewicz 1914; Kohút & Janák 1994; Gawęda et

al. 2016). Crystalline rocks are partly covered by Mesozoic

sedimentary successions as a result of the Alpine orogenesis.

The metamorphic envelope to the polygenetic granitoid

intrusion is exposed mostly in the western part of the massif

and is strongly migmatized at 365–360 Ma (Burda & Gawęda

420

GAWĘDA, SZOPA, CHEW, KLÖTZLI, MÜLLER, SIKORSKA and PYKA

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

a Cameca SX 100 electron microprobe operating in the

wavelength-dispersive spectroscopic (WDS) mode under the

following conditions: 15 kV accelerating voltage, 20 nA beam

current, 1–5 μm beam diameter, peak count-time of 20 s and

background time of 10 s. Standards, analytical lines, dif-

fracting crystals and mean lower detection limits (in wt. %)

were as follows: barite — S (Kα, PET, 0.04), apatite — P

(Kα, PET, 0.03), diopside — Mg (Kα, TAP, 0.02), orthoclase

— Al (Kα, TAP, 0.02), Si — (Kα, TAP, 0.02) and K (Kα,

PET, 0.03), GaAs — As (Mα, PET, 0.04), ThO

— Th (Mα,

PET, 0.04), wollastonite — Ca (Kα, PET, 0.03), Y

— Y

(Lα, TAP, 0.04), La-rich glass — La (Lα, PET, 0.03), CeP

— Ce (Lα, PET, 0.03), NdGaO

— Nd (Lβ, LIF, 0.03), STiO

— Sr (Lα, TAP, 0.03) hematite — Fe (Ka, LIF, 0.09),

rhodohrosite — Mn (Kβ, LIF, 0.03), scapolite — Cl (Kα,

PET, 0.03), phlogopite — F (Kα, TAP, 0.04). The apatite

analyses have been normalized to the sum of 50 negative

charges including 24 oxygen ions and two monovalent anions

(fluorine site), according to the ideal chemical formula of

fluor apatite: A

(BO

)

(X)

where site A is occupied by Ca,

Fe, Mn, Mg, Th, REE, Y and Na, site B by P (substituted by

S, Si) and site X by F, Cl and OH group. For all others mine-

rals, their normalization parameters as well as limits of detec-

tion (LoD) are given in tables.

Cathodoluminescence

CL images of feldspars were obtained using CCL 8200 mk3

apparatus (Cambridge Image Technology Ltd.), mounted on

an Optiphot 2 Nikon microscope in Polish Geological Insti-

tute — National Research Institute, Warsaw. The applied

acceleration voltage was 20 kV, the beam current 500 mA

and the vacuum 67–27 Pa.

Isotopic analyses

The Sm-Nd and Rb-Sr analytical work was performed at

the Laboratory of Geochronology, Department of Litho-

spheric Research, University of Vienna. Results are based on

ID-TIMS procedure.

Sample digestion for Nd-Sr analysis was performed in

Savillex

beakers using an ultrapure 4:1 mixture of HF and

HNO

for 10 days at 110 °C on a hot plate. For whole rock

powders, a minimum dissolution time of 3 weeks was applied

to ensure complete leaching of the REEs from refractory

material. After evaporating the acids, repeated treatment of

the residue using HNO

and 6 N HCl resulted in clear solu-

tions for all samples. The REE fraction was extracted using

50W-X8 (200 – 400 mesh, Bio-Rad) resin and 4.0 N

HCl. Neodymium was separated from the REE fraction using

teflon-coated HdEHP, and 0.24 N HCl as elution media.

Strontium separation followed conventional techniques,

using AG

50W-X8 (200 – 400 mesh, Bio-Rad) resin and

2.5 N HCl as eluants. Maximum total procedural blanks were

< 1 ng for Sr and 50 pg for Nd and were taken as negligible.

Neodymium and strontium were run as metals from a Re

double filament, using a ThermoFinnigan

Triton TIMS,

using La Jolla (Nd) and the NBS987 (Sr) international stan-

dards, respectively. Within-run mass fractionation for Nd and

Sr isotope compositions (IC) was corrected for relative to

146

Nd /

144

Nd = 0.7219, and

Sr /

Sr = 0.1194, respectively.

Uncertainties on the Nd and Sr isotope ratios are quoted

as 2σ

Laser-ablation inductively coupled plasma mass spectro-

metry (LA-ICP-MS) of quartz

Trace-element contents in quartz were determined using

laser ablation inductively coupled plasma mass spectrometer

(LA-ICP-MS) at the Geological Survey of Norway,

Trondheim. It is a double focusing sector field instrument

(ELEMENT-1 Finnigan MAT) combined with a New Wave

UP-193 nm excimer laser probe. Continuous raster ablation

was carried out, resulting in ablated rasters of approximately

150 ×100 nm with depths of 20 to 30 μm. Element concentra-

tions were calculated by multi-standard calibration. Limits of

detection (LoD) are listed in table of data. The analytical

error ranges within 10 % of the absolute concentration of the

element. Detailed description of the measurement procedures

were given by Flem et al. (2002) and Flem & Müller (2012).

Fluorapatite analyses and dating

Apatite crystals separated by handpicking were mounted in

25 mm diameter epoxy resin pucks, then ground and polished

to expose the grain interiors. The fluorapatite crystal mor-

phologies were imaged by scanning electron microscopy on

a FET Philips 30 electron microscope (15 kV and 1 nA) at the

Faculty of Earth Sciences, University of Silesia, Sosnowiec,

Poland.

For trace element and isotopic analyses apatite crystals

were selected from mafic and felsic components. Apatite

U- Pb data were acquired using a Photon Machines Analyte

Exite 193 nm ArF Excimer laser-ablation system coupled to

a Thermo Scientific iCAP Qc at the Department of Geology

Trinity College Dublin. Twenty eight isotopes (

Cl,

Ca,

Mn,

Sr,

139

La,

140

Ce,

141

Pr,

146

Nd,

147

Sm,

153

Eu,

157

Gd,

159

Tb,

163

Dy,

165

Ho,

166

Er,

169

Tm,

172

Yb,

175

Lu,

200

Hg,

204

Pb,

206

Pb,

207

Pb,

208

Pb,

232

Th,

238

U and mass

248

(

232

O) were

acquired using a 50 µm laser spot, a 4 Hz laser repetition

rate and a fluence of 3.31 J/cm

. A Madagascar apatite

(473.5 ± 0.7 Ma; Cochrane et al. 2014) was used as the pri-

mary apatite reference material in this study while McClure

Mountain syenite apatite (523.5 ± 2.1 Ma; Chew & Donelick

2012) was used as a secondary standard. NIST 612 standard

glass was used as the apatite trace element concentration

refe rence material. The raw isotope data were reduced using

the “VizualAge” data reduction scheme (Petrus & Kamber

2012) of the freeware IOLITE package of Paton et al. (2011).

Sample-standard bracketing was applied after the correction

of down-hole fractionation to account for long-term drift in

isotopic or elemental ratios by normalizing all ratios to those

421

AGE AND ORIGIN OF FLUORAPATITE FROM THE TATRA MTS.

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

of the U-Pb reference standards. Common Pb in the primary

apatite standard was corrected using the

207

Pb-based correc-

tion method using a modified version of the VizualAge DRS

(Chew et al. 2014). Over the course of two months of analy-

ses, the McClure Mountain apatite secondary standard

(

207

Pb/

235

U TIMS age of 523.51 ± 1.47 Ma; Schoene &

Bowring 2006) yielded a U-Pb Tera-Wasserburg Concordia

lower intercept age of 524.5 ± 3.7 Ma with an MSWD = 0.72.

The regression line was anchored to a lower intercept using

207

Pb /

206

Pb value of 0.88198, derived from an apatite

ID-TIMS total U-Pb isochron (Schoene & Bowring 2006).

REE contents were normalized to C1 chondrite (Sun &

McDonough 1989).

Results

Petrography and mineral chemistry

Plagioclase crystals in the felsic component of the vein are

andesine (An

– An

), they show typical mechani-

cal twinning (so-called ladder structures; Fig. 3a) and

a greenish CL colour (Fig. 3b). Locally they reveal antiper-

thitic exsolutions of K-feldspar with a composition of Or

93-95

(Table 1). Plagioclase in the mafic component has an oligo-

clase composition (An

– An

). Small

K-feldspar crystals straddle the contact between apatite and

plagioclase crystals (Fig. 4a) and coexist with REE-rich epi-

dote and muscovite (Fig. 4b). Formation of secondary tiny

epidote needles, showing bright yellow-green CL colours

(Fig. 3c; Table 2), is a product of alteration (albitization) of

plagioclase. Mean value of

207

Pb/

206

Pb ratio from 6 pla-

gioclase crystals equals 0.8611 ± 0.0051.

Quartz in the felsic component is intergrown with pla-

gioclase while in the mafic component milky to smoky quartz

fills the wedge-shaped intracrystalline fractures in pla-

gioclases. The trace elements content in quartz differs slightly

between the two components (Table 3). Aluminium, the most

common trace elements in quartz (e.g., Götze et al. 2001;

Müller et al. 2003), is low in both zones, with similar concen-

trations compared to published Al concentrations in mag-

matic and hydrothermal quartz elsewhere (e.g., Jourdan et al.

2009; Müller et al. 2010), while Li is very low, even when

compared to low-Li post-magmatic quartz from the High

Tatra Mts. (Gawęda et al. 2013). Titanium content is higher in

quartz in the mafic component (26.6 – 34.7 ppm) in relation

of quartz in the felsic component (23.4 – 28.3 ppm; Table 3).

Pale-green fluorapatite crystals are present in both compo-

nents and exhibit different morphologies: from ball-shaped

to prismatic ones (Fig. 2 c, d), but no important differences

in chemistry were noted. They are all fluorapatite (Table 4),

rich in REE (total REE content = 2009– 4940 ppm). Yttrium

contents range from 750 ppm to 1468 ppm, while Sr range

from 261 to 638 ppm. High Mn contents (1220 – 2496 ppm)

cause the typical yellow CL emission (Table 5; Fig. 3d). The

chondrite-normalized REE diagram exhibits a minor LREE

enrichment (Ce

/ Yb

= 5.19– 8.17), together with the pre-

sence of a pronounced negative Eu anomaly (Eu / Eu* = 0.24–

0.29) and a slight positive Ce anomaly (Ce /Ce* = 1.09–1.12;

Fig. 5; Table 5). The moderately negative slope of the LREE

patterns (La

/ Sm

= 0.3– 0.6) and strongly fractionated

HREE patterns (Gd

/ Lu

= 11.7–13.0) are relatively consis-

tent throughout the analysed grains (Table 5, Fig. 5). The

Sr /

Sr ratio in apatite is high and equal to 0.719175, with

an age-corrected

143

Nd /

144

Nd ratio of 0.511699 (Table 6).

Fig. 3. Microphotographs and cathodoluminescence (CL) images

showing textural aspects of the apatite-rich pegmatite-like dike

from Baranec Mt. a — ladder structures in plagioclase, crossed

polars; b — CL image of plagioclase with characteristic

yellowish-greenish colour; c — apatite crystal with characteristic

bright yellow CL.

423

AGE AND ORIGIN OF FLUORAPATITE FROM THE TATRA MTS.

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

common Pb. The apatite sample yields a Tera-Wasserburg

lower intercept age of 328.6 ± 2.4 Ma (MSDW = 0.75; Table 11;

Fig. 7). The discordia was anchored using the Stacey &

Kramers (1975) terrestial Pb evolution model calculated for

the apatite U–Pb age. The corresponding unanchored lower

intercept age is 334 ±14 Ma.

Whole rock chemistry of fluorapatite-rich vein

The felsic component shows high SiO

(72.1 wt. %), Al

(16.0 wt. %), Na

O (4.7 wt. %.) and CaO (3.8 wt. %) concen-

trations, while in the mafic component high P

(10.5 wt. %)

and total Fe as Fe

(9.8 wt. %) contents are characteristic

(Table 12). Among trace elements Y and V show the highest

contents in the mafic component. Rare earth element (REE)

concentrations also differ in both vein components (Fig. 5;

Table 12). Chondrite (C1)-normalized REE pattern of the fel-

sic component displays a positive Eu anomaly (Eu / Eu* =

5.45) and weak REE fractionation (Ce

/ Yb

= 4.18). The

REE pattern of the mafic component characterized by a nega-

tive Eu anomaly (Eu / Eu* = 0.28) and also weak REE frac-

tionation (Ce

/ Yb

= 7.09; Fig. 5; Table 12). Isotope compo-

sition of T sample recalculated to 328 Ma is (

143

Nd /

144

Nd)

328

0.511711, ε

328

= – 9.9 and I

328

= 0.72837 (Table 6). Trace

element “spider” diagrams normalized to primitive mantle

(PM) for all samples do not show a significant fractionation

between large-ion lithophile elements (LILE) and high field

strength elements (HFSE) (Fig. 8). Strongly negative Th

anomalies, coexisting with U enrichment, negative Zr and Hf

anomalies and slightly negative Nb anomalies are characte-

ristic of all the samples, while Ta anomalies are lacking.

Posi tive and negative Sr anomalies in the felsic and mafic

components are observed respectively (Fig. 8), similarly to

the positive and negative Eu anomalies in chondrite-normali-

zed REE diagrams (Fig. 5), consistent with plagioclase being

the main carrier of Sr and Eu.

Discussion

Emplacement mechanism and pressure-temperature-

oxygen fugacity estimations

The textures observed in feldspars, like mechanical twin-

ning and the ductile deformations of biotite flakes point out

the vein intrusion was syntectonic (Brown & Parsons 1994).

The cooling time for a relatively small dyke (approx. 1 m in

diameter) should not be long, so one can expect the chilled

margin around the contact in the case of the temperature dif-

ference, which, however, is not present. This suggests that

host rocks were warm enough to prevent the chilled margin

formation or metasomatic character of the vein.

Compound

(wt. %)

LoD

Allanite-(Ce)

Epidote

SiO

0.05 30.96 31.56 32.41 33.17 36.75 37.92

TiO

0.03

0.04

0.16 b.d.l.

0.18

0.03

0.13

0.23

0.18 b.d.l. b.d.l. b.d.l.

0.01 18.14 18.67 20.74 19.70 23.81 26.37

0.05

0.27

0.62

2.31

1.05

0.15

0.05

0.08

3.34

2.96

2.54

2.39

0.16

0.08

0.01

9.60

8.56

6.79

6.49

0.37

0.10

0.01

1.69

1.20

1.34

1.11

0.06

0.10

0.29

6.68

6.07

4.56

3.86 b.d.l.

0.29

0.01

1.68

1.58

0.87

0.00

0.33

1.10

0.76

1.21

0.60 b.d.l. b.d.l.

0.07

1.23

1.94

1.10

4.81 12.03

9.36

FeO

----

10.87

9.82

7.94

6.72

0.09

0.10

MnO

0.05

1.57

1.84

0.52

0.02

0.03

MgO

0.01

0.17

0.14

0.48

0.14

0.11

0.01

CaO

0.03

9.32 10.47 12.87 14.92 22.58 23.46

0.15

0.23

0.16 b.d.l. b.d.l. b.d.l. b.d.l.

O calc.

1.50

1.56

1.63

1.65

1.86

1.89

Total

99.72 99.43 98.38 98.93 98.05 99.80

Crystal-chemical formulae calculated to 8 cations
Si

2.992 2.994 2.989 2.998 2.979 2.988

0.003 0.011

–

0.012 0.002 0.002

0.001 0.005 0.002

–

2.066 2.087 2.254 2.099 2.275 2.448

0.014 0.031 0.113 0.050 0.006 0.002

0.119 0.104 0.086 0.080 0.005 0.002

0.340 0.297 0.229 0.215 0.011 0.003

0.060 0.041 0.045 0.037

–

0.231 0.206 0.150 0.125 0.002 0.003

0.056 0.052 0.028 0.027

–

0.035 0.024 0.037 0.018

–

0.089 0.139 0.077 0.327 0.734 0.555

0.120 0.147 0.041 0.040 0.001 0.002

0.024 0.020 0.066 0.018 0.013 0.001

0.878 0.779 0.613 0.508 0.006 0.007

0.965 1.064 1.271 1.445 1.961 1.980

∑ REE

0.840 0.723 0.576 0.501 0.023 0.014

Note: All iron has been measured as Fe

; b.d.l. — below detection limit.

LoD — limit of detection.

Table 2: Selected analyses of allanite-(Ce) and epidote with their

crystal-chemical formulae (according to Armbruster et al. 2006).

Sample/

Element

LoD

“F”

“M”

0.07

2.20

2.36

1.54

1.33

1.80

1.33

1.04

2.93

2.09

2.61

4.91

3.07

5.41

0.11

0.39

0.53

0.46

bdl

0.34

0.22

0.06

1.34

1.05

1.00

0.78

1.08

1.05

0.01

1.52

4.26

1.28

0.01

bdl

1.48

6.6

43.91 38.71 40.43 44.57 31.20 46.90

2.5

b.d.l.

2.61

2.87

b.d.l.

1.5

24.86 28.34 23.36 26.64 29.55 34.68

min

[

594

606

589

600

609

624

max

[

613

625

608

620

629

645

mean

[

604

616

598

610

619

634

Note: All data obtained by laser ablation inductively coupled plasma mass

spectrometry (LA-ICP-MS). The crystallization temperatures of quartz
T

min

, T

max

, T

mean

were calculated applying the Ti-in-geothermometer by

Wark and Watson (2006). LoD = Limit of Detection. “F”— felsic and

“M”— mafic.

Table 3: Trace elements content of quartz from fluorapatite-rich

dyke and calculated crystallization temperatures.

424

GAWĘDA, SZOPA, CHEW, KLÖTZLI, MÜLLER, SIKORSKA and PYKA

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

The Ti concentrations in quartz (Table 3) have been used

for calculation of the quartz crystallization temperature,

applying the Ti-in-quartz geothermometer of Wark & Watson

(2006). Quartz from the felsic component crystallized

between 598 and 616 °C, while quartz from the mafic com-

ponent exhibits crystallization temperatures 610 – 634 °C

(Table 3), both supporting the late magmatic origin of the

segregation. The presence of antiperthitic exsolution in pla-

gioclase enables ternary feldspar temperatures to be calcu-

lated. Results for the computations yield temperatures of

495 – 550 °C (Table 1), interpreted as cooling conditions.

The amount of phengite substitution in the

muscovite core was used to calibrate the pressure

(Table 7; Massone & Schreyer 1987) which is in

the range of 400 –500 MPa.

The regression equations according to Krani-

diotis & MacLean (1987) and Cathelineau &

Nieva (1985) procedures were used to establish

the temperature range of the chlorite’s formation

from the chemical composition of chlorites

formed at the expense of biotite. The temperature

calculations reveal one crystallization interval,

with the average temperature at 323 °C and at

336 °C, respectively (Table 8; Fig. 9) that could

be consistent with the low pressure muscovite

rims (Table 7), suggesting a clockwise cooling

P -T path. Late fluid circulation causing the

secon dary alterations in this temperature range

and at very low pressure (Fig. 9) is in agreement

with P-T conditions stated elsewhere in the Tatra

Mountains (Gawęda & Włodyka 2012).

The enrichment in LREE and the presence of

pronounced negative Eu anomalies in the chon-

drite-normalized REE patterns of fluorapatite

(Fig. 5), the relatively high Ti content of biotite

(Table 7) and the presence of primary ilmenite

(Fig. 4d), indicate oxidizing conditions in an early crystalli-

zation stage of the mafic component. The presence of minerals

from the epidote group both as inclusions in apatite (Fig. 3c)

and as the late mineral, overgrowing fluorapatite (Fig. 4b),

suggest more oxidizing conditions. The presence of the secon-

dary titanite coronas on ilmenite (Fig. 4d) also associated

with biotite chloritization marks the late, low temperature

oxidation event (Harlow et al. 2006; Fig. 9). The positive Eu

anomaly in the felsic component of a dyke is likely a result of

the abundance of plagioclase which is a carrier of Eu

, sug-

gesting rather reduced conditions in felsic component of the

pegmatite-like segregation. All these facts could be a result

of local oxygen fugacity fluctuations during crystallization.

Timing of the vein intrusion and its relation to the granite

Taking into account the relatively low closure temperature

of the U-Pb system in fluorapatite (375–550 °C; Chamberlain

& Bowring 2000; Schoene & Bowring 2007; Cochrane et al.

2014), and the crystallization temperatures obtained in this

study, the apatite U-Pb age of 328 Ma could be interpreted as

the time of the cooling of the dyke. The apatite U-Pb system

records subsequent cooling of the granitoid intrusion at

340 Ma (Gawęda et al. 2014), and these rocks seem unlikely

to be the source of apatite-rich melt. Fluorapatite-rich rocks

are present in the Tatra Mountains (Gawęda 2008; Szopa et

al. 2013), however they are commonly associated with the

oldest magmatic episode at ca. 368 Ma (Burda et al. 2011;

Szopa et al. 2013) and show U-Pb apatite cooling age of

~ 340 Ma (Gawęda et al. 2014). Multi-stage boron-rich fluids,

also enriched in phosphorus and fluorine, have been

component LoD

0.70

42.63

42.15

41.89

42.46

42.29

42.79

42.25

FeO

0.01

0.12

0.11

0.16

0.33

0.22

0.36

MnO

0.01

0.49

0.70

0.07

1.10

0.55

0.37

0.42

CaO

0.56

54.81

54.54

54.39

54.53

54.88

55.07

54.67

0.01

0.10

0.12

b.d.l.

0.09

b.d.l.

O calc.

0.48

0.35

0.36

0.45

0.06

0.17

0.32

0.17

2.78

3.01

2.96

2.83

3.65

3.45

3.10

0.01

0.00

0.07

0.00

–O=F+Cl

1.17

1.27

1.25

1.21

1.54

1.45

1.31

Total

100.24

99.72

98.54 100.48 100.22 100.61

99.81

Crystal-chemical formulae calculated to 13 anions
P

3.015

3.004

3.013

3.005

3.000

3.016

3.005

0.008

0.011

0.023

0.015

0.025

0.035

0.050

0.005

0.078

0.039

0.026

0.030

4.906

4.919

4.951

4.885

4.927

4.912

4.921

0.016

0.020

0.000

0.015

–

4.965

4.997

4.964

4.989

4.953

4.976

–

0.265

0.199

0.205

0.252

0.033

0.092

0.176

–

0.735

0.801

0.795

0.748

0.967

0.908

0.824

–

0.010

–

1.000

Table 4: Representative electron-microprobe analyses and crystal-chemical for-

mulae of fluorapatite.

Fig. 5. Chondrite (C1)-normalized REE patterns of apatite crystals

from apatite-rich pegmatite-like dyke, whole rock sample (T), mafic

component (M) and felsic component (F).

425

AGE AND ORIGIN OF FLUORAPATITE FROM THE TATRA MTS.

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

Table 5:

Representative LA-ICP-MS analyses of Mn, Sr

, Y

, REE contents in apatite crystals with selected petrological i

ndicators.

Sample

Element (ppm)

Calculated indicators

Eu/Eu*

Ce/Ce*

/Sm

/Yb

/Lu

R1_1

2330.0

589.1

141

1.0

377.4

1270.0

203.1

1088.0

397.9

42.22

506.4

76.25

382.6

52.01

97.5

9.48

41.86

4.82

0.29

1.10

0.59

8.17

12.99

R1_2

19.0

538.4

1381.0

1.4

1108.0

181.8

995.0

373.2

36.94

481.9

73.15

370.5

50.78

97.0

9.29

41.36

4.61

0.26

1.12

0.52

7.22

12.92

R1_3

2271.0

568.9

1437.0

351.7

1220.0

197.3

1061.0

392.5

39.52

491.7

76.34

385.7

53.59

102.0

9.81

44.20

4.9

0.27

1.1

0.56

7.44

12.40

R1_4

2496.0

637.7

1555.0

413.7

1378.0

218.4

1176.0

434.4

46.20

545.6

83.10

416.8

56.88

108.0

10.49

47.16

5.28

0.29

1.10

0.60

7.87

12.77

R1_5

2363.0

539.5

1416.0

355.6

1217.0

195.5

1045.0

382.4

36.91

481.1

74.54

377.4

53.07

101.0

9.76

43.15

4.95

0.26

1.1

0.58

7.60

12.01

R1_6

2433.0

557.9

1440.0

370.9

1241.0

198.8

1060.0

389.3

38.10

492.2

76.30

389.3

53.81

102.7

9.75

44.39

5.16

0.26

1.10

0.60

7.53

11.79

R1_7

2437.0

547.8

1396.0

363.0

1215.0

198.5

1049.0

375.6

37.37

479.1

73.84

375.7

52.55

100.9

9.72

43.77

4.93

0.27

1.09

0.61

7.48

12.01

R1_8

1291.0

261.2

749.7

107.6

469.2

86.9

508.6

202.1

18.12

256.5

40.01

205.2

28.74

54.05

5.28

24.24

2.603

0.24

1.16

0.33

5.21

12.18

R1_9

2181.0

513.2

1353.0

333.9

1157.0

187.7

1017.0

367.9

35.47

468.6

72.40

369.9

51.48

98.0

9.44

42.73

4.77

0.26

1.1

0.57

7.29

12.14

R1_10

1967.0

490.8

1343.0

290.9

1059.0

176.3

973.0

360.9

33.72

460.8

71.47

364.8

51.18

98.0

9.48

42.02

4.74

0.25

1.12

0.51

6.79

12.02

R1_1

2234.0

526.3

1383.0

336.0

1158.0

188.1

1022.0

371.0

36.67

471.1

73.22

373.1

51.80

99.4

9.59

43.18

4.88

0.27

1.10

0.57

7.22

11.93

R1_12

2156.0

1.8

1360.0

319.7

1121.0

183.8

1001.0

368.3

34.53

466.7

71.95

367.8

51.45

98.2

9.56

42.88

4.87

0.25

1.1

0.54

7.04

11.84

R1_13

2329.0

550.7

1406.0

358.7

1216.0

196.4

1035.0

380.9

38.13

482.5

74.31

377.2

52.46

100.6

9.51

43.68

4.97

0.27

1.10

0.59

7.50

12.00

R1_14

2267.0

532.2

1394.0

338.2

1161.0

188.4

1023.0

373.8

36.41

472.3

73.43

372.9

52.31

99.8

9.50

43.45

4.96

0.26

1.10

0.57

7.20

11.77

R1_15

2293.0

539.2

1410.0

351.1

1207.0

194.4

1037.0

379.4

37.18

482.0

74.34

381.0

52.52

100.9

9.64

43.53

4.96

0.26

1.1

0.58

7.47

12.01

R1_16

2287.0

544.5

1422.0

353.5

1212.0

196.3

1050.0

384.7

38.00

485.2

74.90

382.9

52.72

100.8

9.73

43.69

4.94

0.27

1.10

0.58

7.47

12.14

R1_17

2360.0

568.8

1414.0

350.7

1204.0

194.6

1047.0

380.6

39.75

487.3

74.90

382.8

53.02

101.8

9.73

43.35

4.9

0.28

1.10

0.58

7.48

12.29

R1_18

2431.0

599.3

1443.0

376.7

1264.0

204.4

1086.0

392.5

42.48

495.2

76.70

392.3

54.52

103.3

9.81

44.36

4.88

0.29

1.09

0.60

7.68

12.54

R1_19

2413.0

593.1

1418.0

372.9

1253.0

201.0

1065.0

387.3

41.44

488.1

75.20

387.5

53.99

101.4

9.74

43.73

4.86

0.29

1.10

0.60

7.72

12.41

R1_20

2015.0

546.4

1448.0

330.2

1186.0

197.7

1053.0

383.3

37.87

488.5

75.66

388.7

54.13

103.3

9.63

44.58

4.92

0.27

1.1

0.54

7.17

12.27

R1_21

2104.0

537.1

1410.0

322.4

1125.0

185.0

1002.0

371.8

36.36

471.4

73.29

379.7

52.61

99.9

9.56

43.47

4.883

0.26

1.10

0.54

6.97

11.93

R1_22

21550

554.4

1468.0

340.9

1173.0

192.7

1037.0

386.8

37.57

484.0

76.46

390.2

54.61

104.3

10.15

44.75

5.16

0.26

1.10

0.55

7.06

11.59

R1_23

1351.0

481.4

1382.0

186.8

817.0

149.8

876.0

348.2

32.37

457.3

72.00

371.1

52.00

99.1

9.49

42.41

4.73

0.25

1.17

0.34

5.19

11.95

R1_24

1822.0

495.7

1441.0

281.5

1040.0

178.8

978.0

371.6

34.41

479.7

74.67

386.4

53.62

102.7

9.80

44.35

5.06

0.25

1.1

0.47

6.32

11.72

R1_25

1894.0

522.8

1423.0

291.4

1062.0

178.0

984.0

370.4

34.91

475.2

74.60

380.2

53.24

101.8

9.64

43.82

4.96

0.25

1.12

0.49

6.53

11.84

R1_26

1220.0

488.3

1387.0

205.6

865.3.0

156.8

899.7

357.5

32.9

463.4

72.83

375.9

51.88

100.3

9.42

42.87

4.76

0.25

1.16

0.36

5.44

12.03

LoD

0.04

2.85

0.01

0.00

0.01

0.00

0.01

0.00

Sample

Rb [ppm]

[ppm]

Rb/

Sr/

328

Sm [ppm]

Nd [ppm]

147

Sm/

144

143

Nd/

144

328

(Ga)

34.1

363.6

0.267826

0.728370

0.727120

72.63

200.76

0.28680

0.512181

0.000005

–9.85

1.82

–

532.17

–

0.719175

372.47

1006.47

0.223713

0.512179

0.000004

–10.086

1.83

Table 6:

Rb-Sr and Sm-Nd analyses of the whole-rock (T

) and fluorapatite separated grains (ap).

426

GAWĘDA, SZOPA, CHEW, KLÖTZLI, MÜLLER, SIKORSKA and PYKA

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

Compound

(wt. %)

LoD

Ferroan phlogophite

Muscovite

core

mantle

rim

SiO

0.01 36.01

36.17

36.18

36.25

35.71

36.29

47.91

46.5

46.67

TiO

0.02

2.83

2.96

2.97

3.04

2.88

3.05

b.d.l.

0.17

0.22

0.01 17.60

17.46

17.62

17.79

18.09

17.45

33.53

33.16

36.18

0.03

b.d.l.

0.03

0.05

0.06

b.d.l.

0.03

b.d.l.

FeO

0.02 17.82

18.04

17.96

18.07

17.89

18.23

0.49

1.99

1.33

MnO

0.03

0.24

0.22

0.33

0.23

0.24

0.25

0.02

b.d.l.

MgO

0.02 10.87

10.77

10.85

11.00

11.25

10.50

2.01

1.71

0.80

BaO

0.02

0.13

0.46

b.d.l.

0.11

b.d.l.

0.55

0.15

0.28

0.02

0.07

b.d.l.

0.10

0.21

0.09

0.04

0.22

0.25

1.21

0.02

9.44

9.60

9.52

9.23

8.87

9.60

11.13

11.09

9.41

O calc.

3.99

4.01

4.02

4.04

4.00

4.01

4.53

4.46

4.56

Total

99.00

99.71

99.56

99.89

99.19

99.42

100.4

99.5 100.66

Crystal-chemical formulae calculated on the basis of 22 O

5.467

5.466

5.462

5.447

5.398

5.489

6.345

6.258

6.140

0.323

0.336

0.338

0.344

0.327

0.347

–

0.017

0.022

2.533

2.59

2.6

2.61

2.65

2.57

1.65

1.74

1.86

0.617

0.49

0.51

0.5

0.54

0.51

3.58

3.52

3.75

–

0.001

0.002

0.007

–

0.001

–

2.263

2.280

2.268

2.271

2.262

2.306

0.055

0.224

0.146

0.031

0.028

0.042

0.029

0.031

0.032

0.003

–

2.461

2.426

2.443

2.464

2.536

2.367

0.396

0.343

0.158

0.007

0.027

–

0.007

–

0.029

0.008

0.015

0.022

0.006

0.029

0.061

0.027

0.010

0.057

0.064

0.309

1.828

1.851

1.833

1.769

1.710

1.853

1.880

1.905

1.580

#fm

0.482

0.488

0.486

0.483

0.475

0.497

0.122

0.395

0.480

Note: #fm = Fe/(Fe+Mg+Mn). b.d.l. — below detection limit

Compound

(wt. %)

LoD

Sample

#10

SiO

0.03

26.65 25.77 26.11 26.54 25.76 25.95 26.22 34.71 35.13 35.22

TiO

0.05

0.08

0.14

0.06

0.07

0.06

0.07

3.14

3.12

3.23

0.04

18.61 19.74 19.18 15.94 19.66 19.86 19.91 16.76 16.45 16.99

FeO

0.15

27.23 27.03 27.36 37.88 29.36 28.20 28.51 22.30 21.60 21.78

MnO

0.14

0.23

0.22

0.32

0.22

0.49

0.42

0.38

0.33

MgO

0.04

15.05 13.91 14.40

8.35 11.68 12.20 12.74

8.29

8.39

8.28

CaO

0.06

b.d.l.

0.04

b.d.l.

0.04

0.08

0.07

0.05

b.d.l.

0.06

b.d.l.

0.06

b.d.l.

2.06

2.01

1.25

O calc.

11.11 12.97 11.44

9.95 11.04 11.09 11.22 11.75 12.21 11.34

Total

99.03 99.84 99.08 98.96 98.16 98.03 99.17 99.57 99.36 98.53

Crystal-chemical formulae calculated on the basis of 28 O

2–

5.664 5.538 5.580 5.912 5.747 5.599 5.592 6.883 6.973 6.958

2.336 2.462 2.420 2.088 2.253 2.401 2.408 1.117 1.027 1.042

2.326 2.545 2.414 2.098 2.262 2.666 2.609 2.935 2.956 3.049

0.008 0.013 0.023 0.010 0.012 0.010 0.011 0.468 0.466 0.480

0.003 0.053 0.022 0.017 0.019 0.154 0.116 1.299 1.366 1.515

4.837 4.805 4.869 7.040 5.602 4.935 4.969 2.400 2.219 2.083

0.041 0.040 0.058 0.042 0.095 0.090 0.076 0.071 0.064 0.055

4.768 4.456 4.588 2.773 3.986 3.924 4.050 2.451 2.482 2.438

0.014

–

0.033 0.062 0.054 0.038

–

0.033

–

0.033

–

1.042 1.018 0.630

–

16.000 16.000 16.000 16.000 16.000 16.000 16.000 16.000 16.000 16.000

Σ Cations

36.00 35.95 35.99 35.98 35.83 35.83 35.87 34.10 34.03 34.01

Fe/(Fe+Mg)

0.50

0.52

0.72

0.59

0.56

0.60

0.59

0.60

(

314

335

328

275

336

348

344

352

353

(

303

318

313

293

317

324

325

326

331

332

Notes: b.d.l. – below detection limit. LoD – limit of detection.

documented within the Tatra

Massif and are associated with

hydrofracturing, boron metaso-

matism and mo lybdenite crys-

tallization, but are dated at

350 ±1 Ma (Gawęda et al. 2013),

which is roughly 22 Ma older

than the pegmatite dated in this

study and hence should be

excluded as the direct source of

P and F. The chondrite-norma-

lized REE patterns of apatite

differ slightly from those of

Tatra granitoids, showing higher

REE fractionation indices

(Table 5; Fig. 5; cf.: Gawęda et

al. 2016).

The U-Pb apatite age, deter-

mined in this study fits well

with the

Ar-

Ar age range

(330 –300 Ma) obtained for micas

from the High Tatra granite

(Kohút and Sherlock 2003). The

U-Th-Pb ages of secondary

monazite-(Ce) crystals from the

Western Tatra Mountains, also

cluster at ~ 330 Ma (Burda &

Dzierżanowski 2005) and they

are in agreement (within age

uncertainties) of the U-Pb fluor-

apatite age of 328.6 ± 2.4 Ma.

Origin of the parent melt/fluid

to the fluorapatite-rich vein

The analysed vein is unzoned

and does not show any typical

pegmatite internal structure

(e.g. Simmons & Webber 2008;

London 2009). The chemical

composition of both mafic and

felsic components are also not

conclusive. The major element

composition of the felsic com-

ponent is governed by the abun-

dant plagioclase and quartz,

while mafic component chemis-

try is a result of fluorapatite and

ferroan phlogopite concentra-

tions. Among trace elements,

V show increased concentration

Table 7: Micro-chemical analyses

of micas from fluorapatite-rich

dyke and their crystal-chemical

formulae.

Table 8: Representative analyses of chlorite-group minerals and their crystal-chemical formulae and

with Fe

/Fe

and OH calculated assuming full site occupancy. Computed temperatures of crystalli-

zation are according to Kranidiotis & MacLean (1987) and Cathelineau & Nieva (1985)

procedures.

427

AGE AND ORIGIN OF FLUORAPATITE FROM THE TATRA MTS.

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

in titanite, while fluorapatite is a main Y carrier, both phases

showing enrichment in the pegmatite mafic component.

Chondrite-normalized REE-patterns of the felsic component,

with positive Eu anomaly, are typical of feldspar-governed

fractions, while negative Eu anomaly in mafic component

mirrors the geochemical charac teristics of apatite and the

minerals of the epidote group. In the primitive mantle-

normalized multi-element diagrams (Sun & McDonough,

1989) negative Nb and positive Ta anomalies are present,

suggesting atypical arc setting (Thirwall et al. 1994) but

another process (fractional crystallization?) overlapping

the source characteristics. Trace element characteristics of

the felsic component (Sr /Y = 86.63; Table 12) suggest slab

melting and its adakitic sensu lato provenance (Moyen

2009), while mafic component trace element chemistry

shows no such relationships (Table 12). Isotopic signatures

also point out no mantle/lower crustal influence, which is

necessary to characterize the adakitic melts, even those not

connected to slab melting (Macpherson et al. 2006).

As apatite is a Rb-free phase, its Sr composition could be

treated as the direct marker of its origin. The high

Sr/

ratio of apatite (0.719175) and age corrected

Sr/

Sr ratio of

the whole rock of 0.72837 suggest its crustal origin. The

Sm-Nd isotopic composition of the whole-rock vein sample

and fluorapatite are in agreement with the Sr data, also sug-

gesting the crustal origin of the melt parent to the vein. Stable

fluorapatite REE chemistry (Fig. 5) suggests only one melt

source. The ε

328

for apatite-rich segregation (– 9.85; Table 6)

is in the same range as for typical Tatra metapelite

(ε

340

= – 11.32; Gawęda 2009; recalculated value

328

= – 11.13; Table 6)

and from analogous metasedimen-

tary rocks (Kohút et al., 2008). The mean

207

Pb /

206

Pb ratio

of feldspars shows a typically crustal value, supporting the

former suggestions. Possibly dehydration-partial melting of

the upper crust was the source of the melt parent for the

dyke in question and the genetic link to the eclogite boudin

could be excluded, as that shows typical mantle values

(ε

360

= 5.0–6.7; Burda et al. 2015).

However, the computed crystallization temperatures

(598–634 °C) are too low to explain the mafic component

presence, which needs much higher temperatures to crystal-

lize. Low solubility of Fe and Mg in granitic (sensu lato)

melts (Puziewicz & Johannes 1988) implies these elements

should be introduced to pegmatitic systems by aqueous

fluids, circulating throughout the variegated country rocks

and causing the mass transfer from wall rocks (Roda et al.

2004 and references therein).

The main problem is to define the rocks, which were the

source for the mafic components. As Rb-Sr and Sm-Nd iso-

tope systems are very sensitive to cation leaching and are

Table 9: Representative analyses and crystal-chemical formulae of

ilmenite.

Table 10: Representative analyses and crystal-chemical formulae of

titanite.

Fig. 6. Projection of the epidote-group minerals composition on the

(REE+U) versus Al [a.p.f.u.] (after Petrík et al. 1995).

Component (wt. %)

LoD

TiO

0.03

51.04

51.47

49.99

0.03

b.d.l.

0.07

2.10

1.22

3.26

FeO

0.07

39.14

39.56

38.44

0.07

6.52

6.26

MgO

0.02

0.09

0.07

0.09

Total

98.95

98.89

98.04

Crystal-chemical formulae based on 6 O

2–

1.954

1.972

1.929

0.001

–

0.080

0.047

0.126

1.667

1.685

1.650

0.281

0.272

0.007

0.006

0.007

∑ Cations

3.990

3.991

3.984

Notes: b.d.l. — below detection limit. LoD — limit of detection.

Compound (wt. %)

LoD

SiO

0.01

30.66

28.73

TiO

0.03

36.19

39.5

0.01

2.52

0.92

0.07

0.48

0.33

0.04

b.d.l.

0.07

1.11

1.97

MnO

0.07

b.d.l.

0.29

MgO

0.02

0.07

0.05

CaO

0.02

28.48

27.12

0.35

0.61

b.d.l.

O=F

0.28

–

Total

100.17

98.99

Crystal-chemical formulae based on 3 cations
Si

1.000

0.964

0.893

0.995

0.101

0.039

0.009

0.007

0.001

–

0.035

0.057

–

0.009

0.005

0.004

0.991

0.973

Note: b.d.l. — below detection limit. LoD — limit of detection.

428

GAWĘDA, SZOPA, CHEW, KLÖTZLI, MÜLLER, SIKORSKA and PYKA

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

Sample

207

Pb/

235

±2σ

206

Pb/

238

±2σ

238

206

±2σ

207

Pb/

206

±2σ

206

Pb/

238

age

±2σ Ma

207

Pb/

235

age

±2 σ Ma

207

Pb/

206

age

±2σ Ma

207

Pb corr

age

±2σ Ma

tot

(ppm)

tot

(ppm)

tot

(ppm)

R1_1

8.39

0.13

0.1

194

0.0012

0.0022

20.6

0.154317

0.5069

0.0078

727.2

6.7

2274

2780

250

128.56

16.9

97.88

3.25

96.85

R1_2

10.83

0.27

0.1384

0.0019

0.0029

20.8

0.151400

0.5660

0.0120

835.0

11.0

2505

3400

300

121.75

36.8

96.31

2.22

75.50

R1_3

9.46

0.14

0.1292

0.0012

0.0023

23.8

0.137785

0.5322

0.0081

783.2

6.8

2383

2870

280

137.20

28.0

95.48

2.67

84.65

R1_4

9.31

0.13

0.1270

0.0010

0.0022

20.2

0.136400

0.5261

0.0077

770.7

5.9

2367

2850

270

109.74

27.4

130.18

4.27

118.90

R1_5

12.27

0.18

0.1514

0.0014

0.0027

22.4

0.1

17791

0.5846

0.0090

908.8

8.0

2627

2950

310

160.50

28.8

97.54

2.58

68.75

R1_6

12.8

0.28

0.1564

0.0030

0.0038

9.0

0.155350

0.5908

0.0083

936.0

17.0

2670

3190

250

131.05

30.2

100.80

3.51

70.60

R1_7

11.43

0.17

0.1474

0.0015

0.0027

23.0

0.124271

0.5682

0.0091

886.4

8.4

2558

2940

290

1.96

28.5

100.25

3.07

73.90

R1_8

13.48

0.27

0.1630

0.0021

0.0033

0.0

0.124205

0.6000

0.0140

973.0

12.0

2717

3610

330

107.89

35.3

54.18

0.33

34.36

R1_9

10.45

0.17

0.1363

0.0013

0.0025

21.2

0.134570

0.5539

0.0087

824.2

7.3

2473

3170

250

149.51

29.6

91.61

1.81

74.41

R1_10

9.99

0.18

0.1342

0.0015

0.0026

18.6

0.144367

0.5369

0.0085

1.8

8.6

2441

2900

240

110.74

17.9

86.69

1.69

73.07

R1_1

8.90

0.12

0.1240

0.0012

0.0023

20.6

0.149584

0.5192

0.0082

753.7

6.9

2327

2980

220

155.60

34.8

94.08

2.79

87.21

R1_12

10.27

0.15

0.1374

0.0017

0.0027

17.9

0.143018

0.5453

0.0086

829.9

9.7

2458

3040

220

94.07

36.9

93.65

2.35

77.51

R1_13

7.44

0.1

128

0.0010

0.0020

20.5

0.157185

0.4782

0.0070

688.7

6.0

2163

3020

240

43.07

35.1

105.21

3.88

115.20

R1_14

8.16

0.14

0.1

186

0.0010

0.0021

20.2

0.149297

0.5017

0.0088

722.3

5.8

2252

3010

230

55.74

32.4

94.99

2.87

95.32

R1_15

7.53

0.13

0.1

123

0.001

0.0020

19.3

0.158588

0.4852

0.0078

686.3

6.1

2176

2740

280

105.1

34.7

99.35

3.34

108.00

R1_16

7.70

0.12

0.1

150

0.0010

0.0020

21.3

0.151

150

0.4858

0.0076

701.8

5.7

2201

3000

200

106.49

30.9

98.37

3.51

104.00

R1_17

11.43

0.23

0.1468

0.0022

0.0031

20.3

0.143850

0.5680

0.0100

882.0

12.0

2557

3250

260

107.75

17.1

100.77

1.89

74.25

R1_18

9.62

0.18

0.1301

0.0017

0.0026

22.1

0.153610

0.5334

0.0083

788.5

9.9

2400

3150

220

141.25

33.2

104.66

2.92

93.20

R1_19

9.79

0.14

0.1306

0.0014

0.0025

19.9

0.146573

0.5425

0.0075

791.3

8.2

2413

3180

250

129.32

31.1

96.61

2.61

84.50

R1_20

10.57

0.18

0.1372

0.0015

0.0026

22.2

0.138123

0.5597

0.0081

829.0

8.5

2486

3120

220

65.52

42.1

85.87

1.07

67.76

R1_21

11.31

0.16

0.1443

0.0013

0.0026

20.1

0.124865

0.5651

0.0093

868.9

7.1

2550

3280

190

114.25

12.0

89.78

2.57

68.30

R1_22

11.01

0.17

0.1403

0.0015

0.0026

25.8

0.132086

0.5633

0.0092

846.2

8.3

2526

2980

350

124.89

13.8

96.85

2.97

75.90

R1_23

8.90

0.21

0.1

177

0.0017

0.0025

18.1

0.180463

0.5490

0.0150

718.3

9.8

2324

3090

300

105.06

13.0

61.00

0.61

55.99

R1_24

11.63

0.22

0.1467

0.0018

0.0029

32.0

0.134753

0.5667

0.0089

883.3

9.8

2571

3220

210

122.82

13.6

83.59

2.51

62.70

R1_25

11.7

0.22

0.1497

0.0014

0.0027

21.6

0.120481

0.5670

0.01

898.9

8.1

2581

2910

270

127.76

16.3

90.32

1.51

65.30

R1_26

8.14

0.18

0.1

178

0.0015

0.0024

21.4

0.172950

0.4980

0.01

717.8

8.8

2242

3060

210

119.1

13.9

62.29

1.36

63.07

usually mirrored in pegmatites showing

mass transfer from mantle-related coun-

try-rocks (compare: Gawęda 1995), the

eclogite, showing mantle characteristics

could be excluded as the source for all

“mafic” cations. The observations of the

sheared metasedimentary rocks revealed,

that in all of them the replacement of bio-

tite by muscovite occur with pronounced

loss of Fe and Mg (Pyka et al. 2014).

Taking into account the crustal signatures

of the pegmatite-like segregation, proved

by the Rb-Sr, Sm-Nd and Pb isotopic sys-

tems (Table 6), the cation leaching from

sheared metapelites, originally rich in Fe,

Mg, Ca, is the only explanation as the

source of the “mafic” cations. Liberation

of P, Ca, Fe, REE and Y was possible due

to monazite/xenotime decomposition by

-F-H

O – rich fluids, while Th was

trapped in ThSiO

phase (Szopa 2009;

Ondrejka et al. 2012). The ThSiO

phase is

a common remnant in pseudomorphs after

monazite-(Ce) from crystalline rocks of

the Western Tatra Mountains (Szopa 2009)

that explains a negative Th anomaly in

primitive mantle-normalized multi-ele-

ment diagrams (Fig. 8) and enrichment in

P and F in the melt.

The source of fluorine is still unknown,

but the circulation of F-rich fluids was

noted in supra-subduction zones, due to

phengite and /or F-amphibole breakdown

(Sheng et al. 2013). The high activity of

fluorine and phosphorus was noted during

hydrothermal alterations elsewhere in the

crystalline cores of the Central Western

Carpathians (Burda & Dzierżanowski

2005; Szopa 2009; Uher et al 2009;

Gawęda & Włodyka 2012; Ondrejka et

al. 2012).

Fluid mobility was possibly enhanced

by post-collisional uplift and extension.

Decompression of the crystalline rocks

could generate the partial melts, showing

chemistry partly inherited from the melted

metapelitic rocks and from the circulating

(P, F, H

O)-rich fluid which stimulated

melting. The melt temperature was too low

to influence the eclogite tectonically

included in the system, but restricted

leaching cannot be rejected, causing the

oxygen fugacity fluctuations. The accumu-

lation of volatile-rich melts could effi-

ciently stimulate the hydraulic opening of

the fractures. Fluid super-saturated magma

Table 1

Representative LA-ICP-MS U-Pb apatite data for fluorapatite-ric

h dyke from Baranec Mt., Slovakia.

429

AGE AND ORIGIN OF FLUORAPATITE FROM THE TATRA MTS.

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

had been easily moved to these extensional fractures according

to the magma pumping mechanism (Demartis et al. 2011).

The original melt may have had the composition of the felsic

component. The crystallization of abundant plagioclase

causes a decrease in Ca / P ratio and limited crystallization of

apatite in the felsic component (Roda et al. 2004). Phospho-

rus, as incompatible element (Bea et al. 1992) progressively

concentrated in the late fluid, mixed with the external fluid,

rich in F and H

O, carrying the elements fluxed from more

mafic metapelites, squeezed as bubbles and crystallizing the

mafic component of the vein.

Although the main cooling stage is dated ~ 340 Ma using

the same apatite U-Pb method (Gawęda et al. 2014), the

younger pulses of uplift and anataxis cannot be excluded.

Possibly the southern part of the Tatra metamorphic cover

was affected by the multistage uplift, starting from a pre-

intrusive episode, leading to the formation of kyanite-quartz

segregations (Pyka et al. 2013, 2014) through syn-intrusive

(Gawęda et al. 2015), late-intrusive (Gawęda et al. 2014) to

post-intrusive segregation, associated with local and

space-restricted magma /fluid flow.

Conclusions

1. The apatite-rich pegmatite-like dyke was emplaced

syn-tectonically into a fault zone at ca. 600 °C and

400 – 500 MPa. Fluorapatite LA-ICP-MS U–Pb age of

328.6 ± 2.4 Ma is interpreted as the cooling age of the dyke.

2. The Sm-Nd, Rb-Sr and

207

Pb /

206

Pb isotope data suggest

a purely crustal source for the melt/fluid parent to the pegma-

tite-like dyke, and no genetic link to the long-lived Variscan

granitoid magmatism. A possibly source for the apatite-rich

pegmatite dyke was (F, H

O)-rich fluids, generated during

exhumation of the crystalline core of the massif liberating Fe,

Mg, Ca, P, REE from originally mafic metapelites, contai-

ning monazite and xenotime and which could also facilitate

local partial melting.

Compound (wt. %)

“F”

“M”

“T”

SiO

72.1

33.2

55.9

TiO

0.15

1.43

0.53

16.00

14.51

21.33

0.02

0.01

1.46

9.51

3.77

MnO

0.03

0.15

0.05

MgO

0.70

5.65

2.11

CaO

3.80

15.68

5.15

4.71

2.03

7.26

0.62

4.09

0.82

0.04

10.82

1.35

LOI

0.74

1.88

1.75

Total

100.37

98.95

99.98

329.2

297.6

563.60

166

590

235

25.9

224.00

34.10

2.3

17.6

2.4

0.0

0.6

0.0

0.3

12.6

1.70

12.5

19.9

17.4

12.5

21.6

21.10

271

7.5

0.6

0.7

0.4

0.0

3.8

277.1

34.2

1.9

17.7

8.3

0.4

1.6

0.7

2.7

69.8

10.50

4.1

217.1

30.00

0.48

38.08

4.88

3.00

198.7

23.60

0.54

70.53

8.61

0.98

7.78

3.08

0.56

99.95

11.67

0.09

14.29

1.67

0.49

73.33

9.03

0.10

9.70

1.25

0.25

19.04

2.08

0.03

1.78

0.23

0.27

8.43

1.08

0.03

1.01

0.11

ASI

1.050

1.396

1.125

#mg

0.66

0.70

0.69

Rb/Sr

0.08

0.75

0.06

Sr/Y

86.63

1.07

16.48

REE

13.62

829.52

107.79

Eu/Eu**

5.45

0.28

0.94

/Yb

4.18

7.09

7.65

/Yb

1.72

9.81

8.94

Notes: Fe

— total Fe as Fe

; LOI — lost of ignition;

ASI = Al

/(CaO+Na

O+K

O-3.33 P

) in molecular units;

#mg = Mg/(Mg+Fe) (in molecular units); Eu/Eu* = Eu/(√Sm·Gd).

Fig. 7. Tera-Wasserburg concordia diagram anchored through com-

mon Pb for apatite from the apatite-rich pegmatite-like dike from

Baranec Mt.

Table 12: Chemical composition and selected petrological indica-

tors of fluorapatite-rich dyke from Baranec Mt., Western Tatra

Mountains.

430

GAWĘDA, SZOPA, CHEW, KLÖTZLI, MÜLLER, SIKORSKA and PYKA

GEOLOGICA CARPATHICA

, 2016, 67, 5, 417 – 432

3. The high contents of fluorine, phosphorus and water

resulted in a low viscosity and high mobility of the intruding

melt. Rapid crystallization with an accompanying drop in

temperature and pressure resulted in disequilibrium crystalli-

zation, forming the patchy nest-like structure of the dyke and

variations in oxygen fugacity during crystallization.

4. Activation of (F, H

O, P)-rich partial melts and fluids,

was possible due to post-collisional and post-magmatic uplift

of the Tatra Mountains crystalline core in the supra-subduc-

tion zone.

Acknowledgements: Piotr Dzierżanowski PhD. and

Mrs. Lidia Jeżak are thanked for their help during micro-

probe work. The extremely careful editorial corrections by

Pavel Uher and reviews of Adam Pieczka and Milan Novák

significantly improved the manuscript. Evgeny Galuskin is

thanked for help in the mineral formula calculations. This

study was financially supported by the National Science

Centre (NCN) grant 2012/07/B/ST10/04366 (given to AG).

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