PETROGENESIS AND MANTLE DYNAMICS OF PALEOZOIC VOLCANISM IN THE SÔNG ĐÀ STRUCTURE

¹NGUYỄN HOÀNG, ²NGUYỄN ĐẮC LƯ, ²NGUYỄN VĂN CAN

¹Institute of Geological Sciences, Láng Trung, Hà Nội; Now at Geological Survey of Japan, Higashi 1-1-1, Central 7^th, Tsukuba, JAPAN 305-8567; E-mail: hoang-nguyen@aist.go.jp

²Geological Mapping Division of the North, Long Biên, Hà Nội;

Abstract: Paleozoic volcanic activity in association with lithospheric extension and formation of the Sông Đà Structure produced mainly alkaline and sub-alkaline basalts with a subsidiary amount of rhyolite, trachyte and transitional types. The basalts are distinguished by high- and low-Ti types. High-Ti (TiO₂ >1 to 3.5 wt%) type is enriched in highly incompatible as well as light rare earth elements and is also high in FeO* and ratios such as Zr/Y, Nb/Y than the companying low-Ti type (TiO₂ < 1%) is interpreted to be result of melting from fertile asthenospheric material with involvement of mafic veins such as clinopyroxenite and amphibolite, etc.. Whereas, the low-Ti type may be product of melting of the fertile and enriched asthenosphere mixing with refractory and depleted lithospheric mantle as the first rose following lithospheric extension and erosion.

Initial ⁸⁷Sr/⁸⁶Sr, ¹⁴³Nd/¹⁴⁴Nd and ²⁰⁶Pb/²⁰⁴Pb isotopic compositions are obtained based on reported age of 283 Ma ranging from 0.7036 to 0.7090, 0.5119 to 0.5124, and 18.32 to 23.50, respectively. Many values fall in enriched fields relative to CHUR calculated for 283 Ma (⁸⁷Sr/⁸⁶Sr = 0.7041 and ¹⁴³Nd/¹⁴⁴Nd = 0.5123). Isotopic enrichment accompanied by negative anomalies at Nb, Ta, and Zr is interpreted as to be derived from mantle source contaminated by crustal material introduced via plate subduction. We suggest two dynamic models that might be driving force for the volcanic activity: 1)Intraplate lithospheric extension, resulted from long-term plate compression; 2) Back-arc spreading at a plate margin. Such the mantle-lithosphere interaction dynamics should eventually lead to progressive melting without the necessity for a mantle plume to present.

I. INTRODUCTION

For many years the Sông Đà structure has been a subject of intensive study and debate on its geodynamics and related volcanism. Some considered the structure as a Triassic or Mesozoic depression (or rift) [7-9, 38], others believed it was a geosyncline with oceanic ophiolite complexes [40]. However, to date, many researchers have come to an agreement, although with no less controversy, that the structure is an intracontinent rifted zone, started possibly from late Permian and lasted until late Triassic as a result of a prolong compression process related to regional deep faults [7-9, 38].

Paleozoic volcanism occurred widely in the Sông Đà structure zone (SDSZ) at Cẩm Thuỷ, Kim Bôi, Viên Nam, Ba Vì, etc. (Fig. 1). The products are mostly of basalt together with a subsidiary amount of intermediate and acidic types, such as andesite, trachyte and rhyolite. Many detailed studies on the volcanic rocks were conducted over years [3, 7, 8, 31, 36, 37]. Aside from the complexity in temporal and spatial relationships among the lavas, there is a commonly reported feature that low- and high-titanium (low-, hi-Ti) basalt types are recognized within the rift zone regardless of spatial distribution [36, 37, and references therein]. The hi-Ti basalt type, in contrast to the low-Ti type, has higher total alkali and lower MgO contents. Meanwhile, for some low-Ti basalts having very high MgO are called as komatiite [3], which is normally found in Archean lithosphere [2].

Figure 1. Distribution scheme of volcanic localities in the Sông Đà Structure zone

Age of the volcanic eruption is controversial. Radiometric age data are scarce, but Balykin et al [3] reported a Rb-Sr age acquired on clinopyroxene separates (?) from a komatiite-basalt suite in the northwestern side of the rift zone to be 257±7.2 Ma (early Permian) with an initial ⁸⁷Sr/⁸⁶Sr of 0.70299±3. Hoang et al [27] reported an Rb-Sr age of 283 ±21 Ma and ⁸⁷Sr/⁸⁶Sr (initial) at 0.70667±0.000056 for a set of samples ranging from basalt to trachyrhyolite in the Viên Nam and Đồi Bù - Suối Chát area. The latter age is in accordance with late Carboniferrous fossils found in limestone lenses interbedded with basaltic layers in the Hoà Bình - Suối Rút area [32].

We collected a set of samples represented for various rock types in different areas within the structure (Fig. 1) to analyze for chemical and isotopic compositions. The data are reported to discuss the petrogenesis, mantle source, possible involvement of the lithospheric mantle in the context of regional geodynamics.

II. GEOLOGY AND PETROGRAPHY

Volcanic rocks range from mafic to acidic in compositions and are divided into two eruptive phases. The first phase includes mafic (ca. 80%) and sedimentary volcaniclastic products. They are thick lava flows of picro-basalt, basalt, andesitic basalt together with thin lenses of basaltic tuff outcropped in the areas of Suối Chát, Viên Nam, Ba Vì and Đồi Bù. Most basalt layers are greenish due to weathering (chloritization and epidotization). Besides, there are a number of sub-extrusive diabase bodies with thickness ranging from 5 to 10 m, in some case up to 30 m, cutting through earlier eruptives. The second phase is represented by lavas of intermediate to acidic types, such as trachyte, andesite, dacite and rhyolite.

Basalts are mostly aphyritic. Porphyritic type has olivine, clinopyroxene and plagioclase in the phenocryst (up to 15 % vol.). Many of phenocrysts, however, were altered. For example, plagioclase was mostly sericitized, olivine became iddingsite and clinopyroxene was chloritized or epidotized. Basalt textures are microdoleritic or intersertal, whereas structures are massive to porous with pores filled with chlorite, epidote or K-feldspar. Trachyte and trachyrhyolite types are porphyritic with idiomorphic plagioclase, K-feldspar and albite in the phenocryst (up to 25 % vol.). The groundmass contains K-feldspar, amorphous quartz and volcanic glass. Phenocrysts of porphyritic rhyolite are quartz and K-feldspar (up to 15 % vol.), and the groundmass contains micro-quartz and K-feldspar aggregates, the latter are partly kaolinized.

III. ANALYTICAL PROCEDURE

Major and some trace elements were obtained in Việt Nam using an X-ray fluorescence spectrometer; whereas rare earths were obtained using instrumental neutron activation analysis (INAA). Isotopic analyses were conducted at the Geological Survey of Japan (GSJ). Detailed analytical procedure, accuracy and precision are given in [28]. Data are shown in Table 1.

For Sr, Nd and Pb isotope analysis, only fresh rock chips were crushed to pieces of 1-2 mm in size. The chips were washed in HCl 3N in 15 ml Teflon beakers for about 30 minutes then washed ultrasonically in the acid for about an hour, followed by multiple rinses with clean water before being ground in an agate mill. About 50 mg of the powder was dissolved in concentrated HNO₃ and HF, repeated with HNO₃. All the used acids were certified as ultra-clean grade. Elemental extractions were described in [28]. Rb, Sr, Nd and Pb isotope ratios were measured on a multi-collector VG Sector 54 thermal ionization mass spectrometer at GSJ. The ⁸⁷Sr/⁸⁶Sr was normalized to ⁸⁶Sr/⁸⁸Sr = 0.1194 and the ¹⁴³Nd/¹⁴⁴Nd was normalized to ¹⁴⁶Nd/¹⁴⁴Nd = 0.7219. The within-run precision (2s) for ⁸⁷Sr/⁸⁶Sr was ±0.000006 to ±0.000009 and ±0.000007 to ±0.000012 for ¹⁴³Nd/¹⁴⁴Nd. During the period of measurement ⁸⁷Sr/⁸⁶Sr of the NBS 987 Sr standard was 0.710265±0.000008 (1s, n = 28) and ¹⁴³Nd/¹⁴⁴Nd for the JNdi-1 (GSJ) Nd standard [see 20] was 0.512105±0.000005 (1s, n = 19). Lead isotopic compositions were corrected for mass fractionation, and are reported relative to the NBS 981 Pb standard values of (mean, 1s, n = 26) 36.564±0.020, 15.453±0.010, 16.908±0.008 for ²⁰⁸Pb/²⁰⁴Pb, ²⁰⁷Pb/²⁰⁴Pb, and ²⁰⁶Pb/²⁰⁴Pb, respectively. Internal precision of the Pb ratios (2s) is less than 0.01% and total blank is less than 50 pg. The data are shown in Table 2.

IV. ANALYTICAL RESULTS

1. Major element compositions

Based on the analytical results the lavas may be divided into two chemical groups, the first with SiO₂ ranging from 59 to 70 (water %) includes trachyte, andesite, dacite and rhyolite. In this article this group is termed as felsic. The second having SiO₂ content ranging from 44 to 50 (wt%) is termed as mafic. However, because of the large difference in TiO₂ content among the mafic types, following previous researchers [36, 37], we divide the group into low- (TiO₂ < 1%) and high-titanium (TiO₂ > 1 %) subgroups (Table 1, Fig. 2). While the felsic type concentrates in the fields of trachyte, trachyandesite, and rhyolite, hi-Ti mafic type having higher total alkalis (up to 6 wt%) than its low-Ti counterpart falls within the alkaline field (Fig. 2). Moreover, hi-Ti rock type has higher FeO (average 12.5 wt% compared to 8.7 wt%) and lower MgO (average 5.6 wt% compared to 13.6 wt% in the low-Ti type) at a similar SiO₂ content (Table 1). A low-Ti sample that falls in the picro-basalt field has MgO = 24.5 %, Ni = 1297 ppm and Cr = 2922; however FeO is much too low (ca. 9.5 wt%) to be considered as komatiite [2, cf. 3, 37]. We, therefore, classify the rock as a picro-basalt that belongs to the low-Ti mafic type.

Figure 2. Classification of volcanic rocks in terms of SiO₂ vs. total alkalis. Note most of the hi-Ti mafic rocks (circles) plot in alkaline field, whereas low-Ti type (diamonds) plots in subalkaline field. Crosses are felsic rocks.

Table 1. Major and trace element compositions of Paleozoic volcanic rocks of the Sông Đà Structure

Sample	VD1010	VD502	VD1008	VD3010/1	VD3150	VD3046/1	VD19	VD2073	VD3037	VD808	VD3092	VD3094	VD3100
Type	1	1	1	1	1	1	1	1	1	2	2	2	2
SiO2	67.3	68.28	65.82	69.7	69.22	60.44	60.12	58.96	61.16	47.54	47.68	50.3	47.36
TiO2	0.3	0.6	0.5	0.6	0.6	0.9	1.2	1.2	1.1	0.8	0.7	0.7	0.7
Al2O3	13.68	13.25	14.67	13.49	13.08	12.89	15.34	17.38	15	10.6	12.91	12.8	12.4
MnO	2.4	0.81	1.84	3.04	2.88	5.11	7.54	8.31	8.21	2.24	3.19	6.86	3.14
Fe2O3	1.72	3.89	2.8	2.44	2.44	2.44	0.11	0.57	1.8	6.75	6.32	2.59	6.65
FeO	0.12	0.09	0.13	0.09	0.08	0.32	0.27	0.09	0.11	0.24	0.17	0.16	0.16
MgO	0.5	0.81	0.71	0.81	1.91	1.01	0.81	1.01	1.21	14.11	11.59	9.07	11.69
CaO	2.1	1.41	1.4	0.14	0.56	4.35	1.68	0.56	0.42	8.97	7.29	8.69	8.41
Na2O	5	4.71	5.5	3.85	4	4.5	5.38	5.42	5	1.67	2.88	4.21	1.95
K2O	3.75	2.03	2.64	4.48	3.85	3.85	4.54	4.48	4.09	1.13	1.63	0.52	0.33
P2O5	0.24	0.2	0.17	0.24	0.17	0.32	0.41	0.39	0.31	0.36	0.24	0.16	0.24
LOI	2.62	2.48	2.5	1.09	1.08	3.8	2.47	1.53	1.56	3.88	4.7	3.17	5.15
Sum	99.79	98.56	98.91	100.3	100.2	100	100.18	99.9	100.33	98.29	99.78	99.77	98.85
Mg#	34.14	27.07	31.13	37.18	58.26	42.46	92.92	75.96	54.51	78.84	76.58	86.19	75.81
Cr	106	162	114	85	148	112	115	192	135	375.6	1010	675	848
Ni	28	1	24	25	30	54	25	32	62	418	233	246	270
Rb	89	55.4	95	102	69	45	85	92	102	5.2	21	3.6	5
Sr	70	72	190	44	39	73	103	100	83	55.2	130	39	77
Zr	68	420	555	548	522	193	215	189	371	12.6	26	15	21
Y	61	46	64	52	60	30	32	26	44	7.2	14	9	8
Nb	51.3	31	45.2	43.6	12.6	19.5	27.6	22.6	33.1	1.2	2.3	1.5	1.9
Ta	0.96	0.91	1.19	0.96	1.2	0.48	0.75	0.73	0.94	0.14	0.04	0.05	0.04
Hf	1.53	1.2	1.93	1.66	1.71	0.88	1.04	1.14	1.68	0.17	0.17	0.14	0.14
La	107.0	143	172.0	121.0	97.7	67.9	67.9	107.0	107.0	4.73	9.0	3.7	8.8
Ce	204.0	252	285.0	228.0	190.0	122.0	122.0	189.0	201.0	8.82	15.9	8.1	15.9
Nd	92.6	88.7	109.0	97.2	80.7	52.6	53.1	69.6	86.9	4.67	9.0	4.9	6.9
Sm	19.0	18.8	25.1	20.8	22.2	10.8	10.8	13.4	17.6	1.4	2.6	1.4	1.7
Eu	3.9	2.07	4.7	4.4	6.1	3.4	3.6	4.2	6.0	0.51	0.4	0.4	0.5
Yb	8.1	6.6	10.2	8.8	11.5	4.6	5.1	5.4	6.6	1.15	2.0	2.1	1.4
Lu	1.4	0.9	1.8	1.5	1.3	0.7	0.8	0.9	1.2	0.2	0.2	0.2	0.2

Table 1 (continued)

Sample	VD3081	VD3051	VD1006	VD1007	VD3066/1	VD3116/1	VD3026	VD3005	VD3044	VD3127/1	VD2101/1	VD3146	VD3113
Type	2	2	3	3	3	3	3	3	3	3	3	3	3
SiO2	47.04	48.24	46.06	48	47.34	49.4	46.02	49.06	47.88	44.36	46.46	47.86	49.26
TiO2	0.8	0.9	3.0	2.8	1.9	2.4	2.4	2.6	2.9	3.0	3.3	3.4	3.6
Al2O3	10.09	13.6	13.79	12.48	13.47	13.05	13.56	12.72	12.93	14	12.99	12.92	12.59
MnO	2.3	6.22	6.34	5.26	8.65	6.62	6.94	6.42	4.86	7.02	6.58	7.82	7.89
Fe2O3	7.26	4.6	7.22	8.62	5.28	6.68	7.4	7.29	9.27	7.9	7.58	6.54	5.82
FeO	0.17	0.2	0.29	0.25	0.23	0.23	0.27	0.27	0.25	0.27	0.25	0.29	0.21
MgO	15.52	9.57	5.74	5.64	5.14	4.63	5.54	5.04	5.54	7.26	5.95	5.44	4.23
CaO	8.13	10.23	8.41	8.55	8.13	8.41	8.55	7.29	8.41	7.71	8.41	5.89	8.27
Na2O	3.52	2.22	3.2	3.75	2.5	2.92	2.85	2.92	2.03	2.5	3.47	4.25	3.06
K2O	0.48	0.42	2.38	0.67	3.25	2.88	3.47	3.38	1.38	1.63	1.25	1.64	2.78
P2O5	0.29	0.27	0.61	0.58	0.71	0.75	0.72	0.61	0.51	0.58	0.61	0.75	0.8
LOI	4.3	2.8	2.44	1.91	2.74	1.39	1.91	1.79	2.39	3.23	1.78	1.96	0.91
Sum	100.47	99.31	99.72	98.73	99.51	99.73	99.74	99.51	98.43	100.01	98.93	98.97	99.76
Mg#	79.22	78.76	58.63	53.84	63.45	55.27	57.17	55.21	51.59	62.10	58.32	59.73	56.44
Cr	1439	533	185	172	155	170	202	128	59	158	209	201	188
Ni	467	201	96	91	64	66	108	79	69	86	91	59	65
Rb	6	6	39.3	9.2	57	56	49	41	28	30	26	25	53
Sr	102	81	445	348	376	275	258	295	269	443	305	175	291
Zr	21	19	100.8	86	117	130	84	120	121	84	73	129	46
Y	10	12	17.9	19.3	17	21	17	16	18	17	14	21	25
Nb	2.3	2.3	9.9	7.6	16.5	13.5	10.0	12.2	11.3	11.0	3.6	12.4	14.5
Ta	0.1	0.11	0.24	0.22	0.23	0.38	0.12	0.39	0.28	0.24	0.5	0.46	0.44
Hf	0.22	0.24	0.61	0.57	0.43	0.80	0.78	0.97	0.56	0.73	0.83	0.75	0.91
La	10.4	4.4	38.9	30.8	35.8	63.2	38.7	39.1	39.4	26.6	34.9	60.3	71.5
Ce	28.3	13.2	70.8	59	65.8	115.0	79.9	76.6	78.8	53.2	72.2	133.0	130.0
Nd	9.0	8.0	41.6	30.4	32.5	62.7	51.1	41.7	41.4	32.0	47.8	46.2	64.1
Sm	2.3	2.1	8.25	7.1	8.0	11.4	9.9	9.4	9.2	7.4	9.0	11.8	12.5
Eu	0.4	1.1	2.96	2.95	3.5	2.1	3.7	3.6	3.2	3.8	4.0	2.7	2.5
Yb	2.6	2.4	2.64	2.77	3.8	3.1	2.5	3.0	2.5	2.8	2.5	4.0	4.6
Lu	0.3	0.4	0.3	0.3	0.4	0.3	0.4	0.5	0.4	0.3	0.4	0.5	0.5

Note: 1: felsic, 2: low-Ti mafic, 3: high-Ti mafic.

2. Trace element chemistry

The felsic samples have high incompatible elements (Rb, K, Th) and rare earth elements (REE), however, they show relatively low contents of high-field strength elements (HFSE) (Table 1, Fig. 3). Samples of hi-Ti mafic group show much higher incompatible elements compared to those of the low-Ti group. For example, both La and Yb are high in hi-Ti samples, besides, La contents are too high relative to Yb leading to La/Yb ratios are much higher than in the low-Ti samples (Table 1). In general, hi-Ti samples are highly enriched in light REE relative to heavy REE. This feature is easily recognized in Figures 3b-c, expressed by steeper angle from La down to Yb in hi-Ti samples compared to low-Ti samples. In addition, similar to that observed for the felsic group, both hi- and low-Ti samples show relatively strong negative anomaly at some HFSE, especially, Nb and Zr. Aside from the negative anomaly configurations of trace element patterns of the mafic samples are much similar to those of intraplate basalts (for example, oceanic island basalt: OIB) and different from mid-ocean ridge basalt (MORB) or arc lavas [12, 17, 33].

3. Isotopic compositions

Isotopic compositions were corrected for initial values at 283 Ma [27] and shown in Table 2. Corrected results show that initial strontium isotopic ratios vary in a much narrower range from 0.7055 to 0.7065, accompanied by ¹⁴³Nd/¹⁴⁴Nd_i at 0.5124 to 0.5123 and ²⁰⁸Pb/²⁰⁴Pb_i, ²⁰⁷Pb/²⁰⁴Pb_i, ²⁰⁶Pb/²⁰⁴Pb_i, respectively, within 39.43 - 46.91, 15.66 - 15.92, 18.91 - 23.57 (Table 2). Strontium and lead isotopic ratios (initial) of low-Ti samples are lower, and ¹⁴³Nd/¹⁴⁴Nd_i ratios are higher than hi-Ti samples. A low-Ti sample from Bản Tăng shows ⁸⁷Sr/⁸⁶Sr_i at 0.7036 and ¹⁴³Nd/¹⁴⁴Nd_i at 0.51236, falling in the Depleted Mantle field (at 283 Ma). Correlation between Sr and Nd isotopes

Figure 3. Incompatible trace element distribution normalized to primitive mantle for (a) felsic, (b) low-Ti mafic and (c) hi-Ti mafic rocks. Also shown is N-MORB representative for comparison. Negative anomalies at Sr may reflect plagioclase fractionation. Normalizing data from [14]. See text for details.

develops in two trends. One negative covariance, connecting depleted mantle field with enriched continental crust, the second is a nearly straight line where Nd isotopes are nearly constant over a large range of strontium isotopes (Fig. 4a). Lead isotopes meanwhile show positive correlation of ²⁰⁶Pb/²⁰⁴Pb_i against ²⁰⁷Pb/²⁰⁴Pb_i and ²⁰⁸Pb/²⁰⁴Pb_i reflecting uniformly isotopic integration among related U, Th and Pb isotopes. Except for a felsic sample that has very high lead isotopic ratios and plots in a separate field, apart from the field for the rest of the samples (Figs. 4b-c).

Figure 4. Correlation between initial isotopic composition of Sr and Nd (a), ²⁰⁶Pb/²⁰⁴Pb vs. ²⁰⁷Pb/²⁰⁴Pb (b), and ²⁰⁸Pb/²⁰⁴Pb (c). Initial compositions are calculated based on 283 Ma age for the lavas [27]. Depleted mantle (DM) and enriched continental crust (CC) relative to CHUR_283Ma. See text for explanations.

V. DISCUSSION

As described above Paleozoic volcanic rocks in the Sông Đà Structure have undergone post-melting modification processes, including fractional crystallization and weathering alteration. Fractional crystallization is evident by low MgO, average about 5.5 (wt%), much lower than would-be primitive composition [13, 15, 19]. Also, negative anomaly at Sr may reflect fractionation of plagioclase (Fig. 3a). Evidence of the volcanic products having undergone weathering alteration includes primary phenocrysts replaced by secondary minerals. For example, olivine is replaced by iddingsite; clinopyroxene is chloritized or epidotized; while plagioclase is vastly sericitized. Under the impact of alteration processes elements such as Rb, K, Na can be highly mobilized compared to HFSE such as Ti, Zr, Y and Nb [29].

Table 2. Sr, Nd and Pb isotopic composition of Paleozoic volcanic rocks from the Sông Đà Structure

Sample ID	VD550	VD502	VD 1010	VD1008	VD1006	VD1007	VD3005	VD 3044	VD 3066/1	VD 3116/1	VD 3146	VD 3094	VD 3100
Rock type	Basalt	Trachy- rhyolite	Trachy- rhyolite	Trachyte	Hi-Ti basalt	Hi-Ti basalt	Hi-Ti basalt	Hi-Ti basalt	Hi-Ti basalt	Hi-Ti Basalt	Hi-Ti Basalt	low-Ti Basalt	low-Ti Basalt
Location	Kim Bôi	Đồi Bù	Đồi Bù	Đồi Bù	Đồi Bù	Đồi Bù	Đồi Bù	Suối Chát	Suối Chát	Suối Chát	Viên Nam	Bản Tăng	Bản Tăng
Rb (ppm)		55.4	89.6	95.2	39.3	9.2	40.8	28.4	57.3	56.2	25.2	3.6	4.8
Sr (ppm)		72	70.2	271.97	402.222	348.1	476.4814	407.4009	540.11	351.42	175	39.2	77.4
^87Rb/^86Sr		2.0915	3.4695	0.9515	0.2656	0.0718	0.2327	0.1895	0.2884	0.4347	0.3914	0.2496	0.1686
^87Sr/^86Sr	0.706857	0.715604	0.720642	0.710311	0.708142	0.706758	0.706617	0.706099	0.707635	0.706991	0.706409	0.704707	0.709512
^87Sr/^86Sri		0.706527	0.705585	0.706182	0.706989	0.706446	0.705607	0.705277	0.706383	0.705104	0.704710	0.703624	0.708780
Sm (ppm)		18.8	19	25.1	8.25	7.12	9.42	9.22	7.96	11.4	11.8	1.36	1.67
Nd (ppm)		88.7	92.6	109	41.6	30.4	41.7	41.4	32.5	62.7	46.2	4.89	6.86
^147Sm/^144Nd		0.1280	0.1239	0.1391	0.1198	0.1415	0.1364	0.1345	0.1479	0.1098	0.1543	0.1680	0.1470
^143Nd/^144Nd	0.512607	0.512559	0.512633	0.512566	0.512561	0.512565	0.512597	0.512475	0.512326		0.512585	0.512670	0.512224
^143Nd/^144Ndi		0.512322	0.512403	0.512308	0.512339	0.512303	0.512344	0.512225	0.512051		0.512299	0.512359	0.511952
U (ppm)		3.22	4.08	8.98	0.31	0.76	0.32	1	0.21	0.28	1.12	0.12	0.1
Th (ppm)		13.8	27.9	27.6	5.21	3.95	5.4	6.58	4.93	4.54	5.07	0.33	1.34
Pb (ppm)		18.4	21.7	65.3	15.1	17.4	20	12.1	21.6	18.1	19.5	13.2	53.1
^238U/^204Pb		11.092	11.917	8.716	1.301	2.768	1.014	5.238	0.616	0.980	3.640	0.576	0.119
^235U/^204Pb		0.080	0.086	0.063	0.009	0.020	0.007	0.038	0.004	0.007	0.026	0.004	0.001
^232Th/^204Pb		49.350	84.600	27.811	22.703	14.937	17.766	35.782	15.018	16.505	17.108	1.645	1.660
^206Pb/^204Pb	19.294	25.863	24.112	19.299	18.545	18.781	18.875	18.812	19.300	18.952	19.035		19.761
^207Pb/^204Pb	15.619	15.942	15.907	15.682	15.647	15.583	15.690	15.571	15.654	15.590	15.646		15.772
^208Pb/^204Pb	39.681	47.608	45.636	39.831	39.034	39.197	39.340	39.171	39.879	39.710	39.361		40.238
^206Pb/^204Pbi		25.366	23.577	18.908	18.487	18.657	18.830	18.577	19.273	18.908	18.872		19.756
^207Pb/^204Pbi		15.916	15.880	15.661	15.644	15.576	15.688	15.559	15.653	15.588	15.637		15.772
^208Pb/^204Pbi		46.912	44.443	39.439	38.714	38.986	39.089	38.667	39.667	39.477	39.120		40.215

(i) Initial isotopic compositions calculated based on 283 Ma reported for the volcanic rocks [29]

1. Crustal contamination

Magma passing through or incubating in the crust may interact with surrounding materials. Crustal contamination of mafic magmas is expressed by having low Nb, Ta, Nb/Y and Ta/Y, and high concentration of incompatible elements, such as Rb, Ba, K, relative to other trace elements [6, 12, 16]. Negative anomalies at HFSE suggest that volcanic rocks in the Sông Đà Structure might be contaminated. Crustal contamination is resulted in having high strontium and low neodymium isotopic ratios that head toward the continental crust field (Fig. 4a). Except for a sample with low Sr and high Nd isotope that falls in the depleted field all the rest of the samples distribute in enriched fields relative to chondrite calculated for 283 Ma (CHUR_{283 Ma:} ⁸⁷Sr/⁸⁶Sr = 0.7041, ¹⁴³Nd/¹⁴⁴Nd = 0.5123). In short, most of the studied samples were contaminated by crustal materials either after or before melting (products of contaminated source mantle).

Figure 5. Plots of ¹⁴³Nd/¹⁴⁴Nd_i vs. Ce/Pb: a) suggest possible crustal contamination for some Sông Đà volcanic rocks contradicting to broadly positive correlation between ⁸⁷Sr/⁸⁶Sr_i vs. Nb/Y; b) Labels as in Fig. 3.

Lead concentration in the crust is high but Nd isotope is low. Interaction between mantle-derived magmas with crustal material results in positive correlation between Nd isotopes and Ce/Pb ratios as illustrated in Fig. 5a [4, 5]. Similarly, effect of crustal contamination may reflect in negative covariance between Sr isotopes and Nb/Y ratios; however, the correlation is positive for most of the samples (Fig. 5b). Therefore, timing of the crustal contamination is inconclusive. Negative and positive correlation between, respectively, Nd and Sr isotopes, and Ce/Pb (Figs. 4a and 5a) may reflect mixing between an enriched and a depleted source. For example, an alkaline rhyolite (Đồi Bù, VD502), traditionally viewed as a crust-derived volcanic rock for usually having low Nd isotope and Ce/Pb (or high Sr isotope and low Nb/Y), shows high Nd isotope and Ce/Pb, plotting closer to the depleted field (Fig. 5b). Another example, sample VD3094 (low-Ti basalt, Bản Tăng) while having high Nd and low Sr isotope distributing within the depleted field (Fig. 4a) has low, crustal-like Ce/Pb ratio (Fig. 5a). Therefore, effect of crustal contamination is inconclusive. Instead, trace element and isotopic characteristics of Permian volcanic rocks in the Sông Đà structure may possibly reflect mixing between depleted and enriched sources.

2. Melting process

Correlation between TiO₂/Al₂O₃ and TiO₂ is not affected by fractionation of olivine and clinopyroxene. Besides, ratios such as Ti/Y, Zr/Y and Nb/Y may be used as indicators to evaluate source mantle and melting process because they are not affected by fractional crystallization and relatively resistant to alteration processes [28, 29, 39]. However, in order to understand mantle source region and melt segregation conditions it is essential that chemical composition of the primitive melt be known.

Volcanic rocks in the Sông Đà Structure have high contents of incompatible elements; and high light-REE relative to middle and heavy REE (Table 1, Figs. 3a-c). Elements such as Sm, Yb and Y are higher in the hi-Ti samples but much lower in the low-Ti lavas relative to MORB [e.g. 14, 15, 33]. The elements are highly compatible in garnet-bearing rocks while incompatible in the presence of other minerals, such as spinel [15, 18, 24]. Mid-oceanic ridge basalts are believed to be derived by melting of refractory spinel peridotite that underwent previous melting events occurred during processes of oceanic crust formation [15, 23, 39, and references therein]. Very lower concentrations of Sm, Yb and Y in the low-Ti mafic samples compared to MORB may suggest that garnet was a residual mineral. In contrast, high contents of the elements in the hi-Ti mafic samples may indicate the lavas were products of spinel peridotite melting. Moreover, higher ratios of La/Yb in the hi-Ti samples relative to low-Ti samples may suggest that they were derived from a more enriched source and/or smaller melting degrees [15, 18, 23, 33]. In short, distribution configuration of rare earth pattern not only reflects geochemical characteristics of source region, but also melt segregation condition and nature of primary magma.

Chemical compositions of primary magmas in the Sông Đà Structure are unknown. However, they can be approached by addition (or extraction) of olivine (and clinopyroxene) that crystallized (or cumulated) during the process of magma evolution to the present composition [29, 41]. As described above, many of the Sông Đà volcanic rocks have olivine and a minor amount of clinopyroxene in the phenocryst. Besides, some cumulate olivine aggregates were observed in low-Ti samples [3, 36, 37]. Therefore, we may conclude that most of the mafic rocks underwent olivine fractionation at early stages and olivine + clinopyroxene at later stages before the magmas erupted to the surface. In order to simplify the calculation task while protecting the meaning of searched primary chemistries, in this article we conduct only olivine correction. The calibration is based on the following criteria: first, primary mafic melt, product of spinel or garnet peridotite melting, has a Mg-number [100*Mg/(Mg + Fe²⁺)] in the range of 69 to 71 [13, 18, 19, 41]; second, olivine - melt Fe²⁺/Mg distribution coefficient (Kd ^Fe++/Mg (Ol/liq) = 0.30 [34]) and composition of olivine is Fo₈₇ to Fo₈₉; third, in order to avoid plagioclase fractionation effect only samples with MgO higher than 6 (wt%) were applied for correction [39]. Based on the above criteria we added Fo₈₅ olivine step by step to mafic rocks at the ratio of 1:99 until the rock reached a Mg-number between 69 to 70.5 and correspondent olivine fell within Fo₈₇ to Fo₈₉. In short, depending on chemical composition, olivine was added (+) or subtracted (-) [41]. Results are shown in Table 3. Note that although the calculated results might not reflect the true chemistry of primary melts they show, however, that fractional crystallization of olivine (and clinopyroxene) increases TiO₂ in the melt while TiO₂/Al₂O₃ ratios are nearly unchanged. As illustrated in Figure 6 that high TiO₂ contents in hi-Ti samples are not derived by melting of either spinel or garnet peridotite (line 2 and 3). Results of experimental melting of peridotite shows that the lower melting degree the higher Ti content in the melt, and that at the same melting parameters the more fertile peridotite produces melt with higher Ti (line 2 compared to line 3 in Fig. 6). In addition, with melting degree smaller than 1%, melting of a fertile peridotite produces melt with maximal TiO₂ 1.5 (wt%) and corresponding TiO₂/Al₂O₃ < 0.1 [9, 15, 23]. Obviously, the hi-Ti mafic rocks, unlike their companying low-Ti, must be derived from a much higher Ti source. Hi-Ti mantle-derived rocks may include clinopyroxenite, wehrlite, websterite, amphibolite, etc. They usually have high TiO₂, Al₂O₃, TiO₂/Al₂O₃ and incompatible elements compared to lherzolite or harzburgite. They normally include hi-Ti clinopyroxene and amphibole (kaersutite), ± phlogopite and secondary minerals such as apatite, rutile, and ilmenite [5, 24]. Experimental melting of mixture of a peridotite and a composite basalt component by Kogiso et al. [18] produced melts with high FeO, TiO₂, and TiO₂/Al₂O₃, similar to the hi-Ti mafic samples reported here (line 1 in Fig. 6). This observation suggests involvement of mafic component(s) in the petrogenesis of hi-Ti mafic lavas in the Sông Đà rift zone.

Hi-Ti lavas have much higher Zr/Y, Ti/Y and Nb/Y compared to low-Ti samples (Figs. 7a, b) Because the correlation between Nb/Y and Zr/Y (Ti/Y) is not effected by low-pressure fractional crystallization (pyroxene ± plagioclase) difference in these ratios in low- and hi-Ti reflects melting degree, depth of melt segregation and source fertility [13, 23]. Besides, melt-solid distribution coefficients (K_d ^liq/solid) of Nb <Zr <Ti <Y [5, 14, 16] high ratios of Nb/Y, Zr/Y and Ti/Y in hi-Ti lavas may suggest that they were derived by lower melting degree and/or from a more fertile source relative to low-Ti magmas.

Figure 6. Relationship between primitive TiO₂/Al₂O₃ vs. TiO₂ for low- and hi-Ti Sông Đà mafic rocks (Table 3). Results of experimental melting [13] for refractory (line 3), fertile peridotite (line 2) and mixed peridotite - mafic component (line 1) [18], arrows point in the directions of progressive partial melting. Distribution field of average world-wide basalts (contoured) is shown for comparison. Note that fractionation of olivine and clinopyroxene does not affect TiO₂/Al₂O₃ ratios.

3. Source characteristics and geodynamics of the Sông Đà volcanism

Above we have shown that melting of a mixture of peridotite and mafic component may produce melts with high Ti, Fe and incompatible elements. Distribution of mafic veins in ultramafic bodies is common [30]. These ultramafic bodies are residues, cumulated after melting of asthenosphere to be the major component of lithospheric mantle. The existence of mafic veins is explained as small volumes of trapped basaltic melt, products of local, discrete melting within the lithospheric mantle [6], or products of metasomatic processes under the influence of asthenospheric heat [24, 25].

Table 3. Primary melt composition calculated using olivine addition for representative samples

Sample	VD3092	VD3094	VD3100	VD3081	VD3051	VD3116/1	VD3005	VD3044	VD3127/1	VD2101/1
% Ol (+/-)	4	11	4	1	14	31	31	31	31	31
SiO2	47.20	49.08	47.40	46.53	46.97	46.03	45.77	45.26	42.61	44.40
TiO2	0.67	0.62	0.68	0.79	0.78	1.60	1.73	1.95	2.10	2.26
Al2O3	12.37	11.45	12.02	9.90	11.80	8.69	8.47	8.71	9.77	8.89
MnO	9.45	9.46	9.83	9.45	10.66	13.15	13.29	13.49	13.70	13.47
FeO	0.16	0.14	0.16	0.17	0.17	0.16	0.18	0.16	0.19	0.17
MgO	12.92	12.92	13.13	15.70	14.31	17.65	17.93	18.34	18.29	18.24
CaO	6.99	7.77	8.15	7.97	8.88	5.60	4.86	5.66	5.39	5.76
Na2O	2.76	3.77	1.89	3.46	1.92	1.95	1.95	1.37	1.74	2.37
K2O	1.57	0.47	0.32	0.47	0.37	1.92	2.26	0.93	1.13	0.85
P2O5	0.23	0.14	0.23	0.28	0.23	0.50	0.41	0.35	0.40	0.42
TiO2/Al2O3	0.05	0.05	0.06	0.08	0.07	0.18	0.20	0.22	0.21	0.25
Mg#	70.91	70.89	70.43	70.15	70.53	70.53	70.65	70.79	70.42	70.72
Fo (Ol)	88.23	89.17	88.28	88.19	88.52	88.24	88.25	88.62	88.49	88.43

Olivine used in the calculation SiO₂: 40.01, FeO: 14.35, MgO: 45.64 (wt%), correspondent Fo₈₅

In general, the lithospheric mantle is composed mostly of refractory peridotite rocks. They are residues after melting of asthenospheric material to form mafic melts, therefore, they are enriched in Mg and depleted in basaltic components such as Fe, Ti, Ca, Al, Na, etc., light-REE and especially, very depleted in highly incompatible elements [1]. However, due to metasomatic activities, interaction with mafic melts introduced from below, trace element budget in the lithospheric mantle may be increased [1, 6, 24, 25]. This re-enrichment process is obviously depended on geochemical environment and duration. In summary, geochemistry of the lithospheric mantle is considered to be complicatedly heterogeneous, and the heterogeneity can be in a large as well as small scale [25]. In contrast, asthenosphere material is believed to be homogenous, fertile in basaltic component, and enriched in trace element concentration [1].

The Sông Đà structure has been considered as a depression (rift) that was developed from late Permian till late Triassic [7-9, 31, 38]. Regardless of whether the rift initiated in an intraplate or marginal environment, during lithospheric subsidence and extension processes, base of the lithospheric mantle will rise and be eroded under heat impact from the asthenosphere that rises up to fill in gaps left behind by lithospheric mantle in accordance with the uniform stretching model [20, 23]. This

Figure 7. Correlation of Nb/Y vs. Zr/Y for Sông Đà Paleozoic low- and hi-Ti mafic rock types relative to fields of mantle-derived primitive basalts (dashed lines) and depleted MORB. The higher Nb/Y and Zr/Y the more enriched source and/or the lower melting degree. See text for details. Labels as in Fig. 3.

interaction leads not only to discrete metasomatism and localized melting to develop within the lithospheric mantle, but eventually ignites decompression melting of the asthenosphere [1, 15, 23]. In this case, mixing between asthenospheric melts with mafic veins should generate hi-Ti volcanic rocks with high Fe and trace elements as we argued above. Whereas, melting of the asthenosphere with a contribution of (garnet-bearing peridotite) lithospheric mantle may produce melts with low Fe/Mg, relatively low trace element concentrations similar to the low-Ti mafic rocks [1, 24, 33, 39]. Difference in the chemical compositions between lo- and hi-Ti mafic types may reflect the difference in participating ratio between mafic veins and/or lithospheric mantle versus asthenospheric material, and degrees of partial melting.

Involvement of mafic components and lithospheric mantle has been invoked to explain the formation of lo- and hi-Ti volcanic rocks in the Oslo rift [26] and the Paranã volcanic province [22]. Participation of clinopyroxenite veins in the melting of asthenosphere is viewed to be essential in the formation of Hawaiian hi-Ti alkaline basalts [18].

Lithospheric mantle is considered to be chemically heterogeneous, “cold”, “dry”, depleted in basaltic component, and low in viscosity [1]. Therefore, it has lower chance of melting to form large volumes of basaltic melts at rifted zones. However, there are opinions that lithospheric mantle can melt to produce basalts if it becomes more viscous [29, 39]. Dehydration of hydrous minerals such as phlogopite, amphibole, etc., may lead to decrease solidus temperature, on the one hand, and raise the viscosity, on the other. In both cases, physical properties of the lithospheric mantle become more asthenospheric, which are more ready to melting [29, 39].

For a long time the Sông Đà Structure has been coined as a Permian-Triassic “intra-continental” rift (depression). We believe “continental” may cause confusion since during the Permian, the territory was possibly oceanic [10, 11, 32, 35]. The volcanism was not an island arc related for there is very high percentage of mafic rocks (80 %) over intermediate and acidic types. There is no evidence of being mid-oceanic ridge volcanism for the latter has no acidic volcanic lavas. Therefore, the environment for the Sông Đà volcanism to appear was most likely intraplate or marginal with an oceanic (not continental) lithosphere that was composed of mafic and ultramafic layers. This argument points out that the crustal contamination discussed above could not have happened during the magma passage to the surface but rather it may happen in the mantle before the melting occurred. Crustal material, especially hydrous minerals in oceanic sediments, may be introduced into the mantle at subduction zones. Under high temperature and pressure condition hydrous minerals dehydrate to form hydrous fluids that carry highly soluble elements such as Ba, Rb, K, Th, etc., to shallower mantle, leaving behind poorly soluble HFSE, such as Ti, Nb, Zr, etc. [5, 12, 16, 17]. Melting of mantle mixing with lesser than 1 % of crustal material produces anomalies not only in Sr, Nd and Pb isotopic composition, but also in the configuration of trace element patterns [see 28 for calculation example]. Mantle may be contaminated by crustal material in global scale (for example, Indian Ocean asthenosphere [21]), or regional scale, for example, mantle in the proximity of West Pacific subduction zone [12, 17, 41]. Ancient subduction zones was widely developed behind Eurasia in accordance with the development and disappearance of the PaleoTethys sea (ca. 400 Ma to 200 Ma [35]). Crustal contamination of source region of the Sông Đà volcanism may be regional that developed over a long period, or it may reflect a local phenomenon for being in proximity to a subduction zone.

From the above discussion we present two geodynamic models for the Sông Đà volcanism: 1) Intraplate lithospheric extension, and 2) Back-arc spreading resulted from subduction. The first model may be viewed as a result of prolong compression related to regional deep fault systems (see above). For the second, extension occurred following erosion of lithospheric base by mantle convection caused by a subduction zone. However, the second model requires that the Sông Đà Structure was located at an active margin.

VI. CONCLUSIONS

1. Permian volcanism in the Sông Đà Structure consists mainly of mafic rocks with a subsidiary amount of intermediate and acidic lavas. The mafic type is classified into hi-Ti (TiO₂ >1 wt%) and low-Ti (TiO₂ <1 wt%). Hi-Ti samples have higher FeO*, total alkali and concentration of trace elements than low-Ti lavas.

2. Hi-Ti magmas are explained by melting of fertile asthenosphere mixing with mafic veins in the lithospheric mantle enriched in Ti and incompatible elements. Whereas low-Ti rocks are explained by mixing between fertile asthenosphere and refractory lithospheric mantle melts.

3. High Sr, Pb and low Nd initial (283 Ma) isotopic composition together with negative anomalies at high-field strength elements (Nb, Ta, Zr, Y) suggested that source mantle might be contaminated by crustal material introduced into the mantle via a subduction zone.

4. Two dynamic models might be driving force for the volcanic activity: 1) Intraplate lithospheric extension, resulted from prolonged plate compression; 2) Back-arc spreading at a plate margin. Both of the dynamics might be related to the development and disappearance of PaleoTethys (ca. 400 Ma to 200 Ma).

5. Mantle-lithosphere interaction dynamics provided by the models should eventually lead to progressive melting to form the volcanic rocks without the necessity for a mantle plume to present.

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