币圈fib什么意思 币圈fil
『壹』 四氧化三铁、氧化亚铁、氧化铁分别和空气经什么条件形成的
四氧化三铁是铁在空气中剧烈燃烧形成的，氧化铁是铁在空气中自我氧化的结果，至于氧化亚铁应该是在空气不足的条件下自我氧化生成的吧（这个我不确定）。
『贰』 铁与氧气在高温下反应成什么
铁和氧气在高温下生成四氧化三铁。在高温时，铁在纯氧中燃烧，剧烈反应，火星四射，生成四氧化三铁。化学方程式如下：
(2)币圈feo是什么意思扩展阅读铁的用途：
1、用于制药、农药、粉末冶金、热氢发生器、凝胶推进剂、燃烧活性剂、催化剂、水清洁吸附剂、烧结活性剂、粉末冶金制品、各种机械零部件制品、硬质合金材料制品等。
2、纯铁用于制发电机和电动机的铁芯,还原铁粉用于粉末冶金,钢铁用于制造机器和工具。此外,铁及其化合物还用于制磁铁、药物、墨水、颜料、磨料等。
3、用作还原剂。用于铁盐制备。还用于制备电子元器件。
4、用作营养增补剂(铁质强化剂)。
『叁』 FeO,Fe2O3,Fe3O4如何区分它们啊！它们有什么联系吗速！谢啦！
氧化亚铁(FeO)是一种黑色粉末，不稳定，在空气里加热，可被氧化成四氧化三铁，和酸(弱氧化性酸)反应。
(3)币圈feo是什么意思扩展阅读：
主要用途：
1、FeO：
可以被用作色素，在化妆品和刺青墨水中都有应用.氧化亚铁也应用于瓷器制作中使釉呈绿色。但是这个物质不稳定，很容易被氧化成四氧化三铁。
2、氧化亚铁：
用于油漆、橡胶、塑料、建筑等的着色，是无机颜料，在涂料工业中用作防锈颜料。用作橡胶、人造大理石、地面水磨石的着色剂，塑料、石棉、人造革、皮革揩光浆等的着色剂和填充剂，精密仪器、光学玻璃的抛光剂及制造磁性材料铁氧体元件的原料等。
用于电子工业、通讯整机、电视机、计算机等磁性原料及行输出变压器、开关电源及其高U及高UQ等的铁氧体磁芯
用作分析试剂、催化剂和抛光剂，也用于颜料的配料；
用于各类药片、药丸的外衣糖衣着色用
用作磁性材料、颜料及制取还原剂、抛光剂、催化剂等；用于药片糖衣和胶囊等的着色
用作防锈漆的颜料。因该品制成的云母氧化铁防锈漆抗水渗性好，防锈性能优异，可以取代红丹
食用红色素。日本用于赤豆饭、魔芋粉食品。对曾用防腐剂处理果柄切口的香蕉加以识别时用。美国多用于猫食、狗食和包装材料
无机红色颜料主要用于硬币的透明着色，也用于油漆、油墨和塑料的着色
广泛用于油漆、橡胶、塑料化妆品、建筑精磨材料、精密五金仪器、光学玻璃、搪瓷、文教用品、皮革、磁性合金和高级合金钢的着色；主要用作磁性材料、颜料、擦光剂、催化剂等，还用于电讯、仪表工业；主要用作磁性材料、颜料、擦光剂、催化剂等，还用于电讯、仪表工业无机红色颜料。
3、四氧化三铁：
四氧化三铁是一种常用的磁性材料。
特制的纯净四氧化三铁用来作录音磁带和电讯器材的原材料。
天然的磁铁矿是炼铁的原料。
用于制底漆和面漆。
四氧化三铁是生产铁触媒（一种催化剂）的主要原料。
它的硬度很大，可以作磨料。已广泛应用于汽车制动领域，如：刹车片、刹车蹄等。
四氧化三铁在国内焊接材料领域已得到认可，用于电焊条、焊丝的生产尚属起步阶段，市场前景十分广阔。
四氧化三铁因其比重大，磁性强的特点，在污水处理方面表现出了良好的性能。
四氧化三铁还可做颜料和抛光剂。
我们还可以通过某些化学反应，比如使用亚硝酸钠等等，使钢铁表面生成一层致密的四氧化三铁，用来防止或减慢钢铁的锈蚀，例如枪械、锯条等表面的发蓝、发黑。俗称“烤蓝”。
『肆』 贺根山式蛇绿岩型铬矿
贺根山式铬矿：产于二连浩特-贺根山-黑河蛇绿岩带，即西伯利亚板块与锡林浩特-松嫩微陆块（属广义的中朝板块）碰撞带，蛇绿岩形成时间为晚泥盆世，侵位时间大约为晚石炭世—早二叠世。岩体由高铝地幔橄榄岩和堆晶岩组成，局熔改造型豆荚状铬矿赋存于中部纯橄岩-斜辉辉橄岩杂岩相中，矿石工业类型以耐火级为主，矿床规模达到中型。堆晶岩相中的结晶分异铬矿，工业价值不大。
一、内蒙古锡林浩特市贺根山铬矿床
（一）概况
贺根山铬矿床位于内蒙古锡林浩特市潮克乌拉苏木乡北西9km，地理坐标：东经116°18＇00〞，北纬45°50＇00〞。该矿床包括3756、620和733三个矿区，其中以赫格敖拉3756矿区规模最大，由186个矿体群组成，截止1993年底，累计探明C+D级储量129.9万吨，属中型铬矿矿床，也是华北地区最大的铬矿床。矿石Cr2O3含量平均为23.63%～27.26%，铬铁比值Cr2O3/＜FeO＞=2.18～2.5。矿体主要为富铝型耐火级铬矿，仅有个别矿体为富铬型冶金级铬矿。
1954年东北地质局126队在该区普查找矿发现超基性岩体及铬、镍等多处矿点，1956年该队在3756探井中首次见矿，1963年提交了《内蒙锡林郭勒盟赫格敖拉3756铬矿床最终勘探报告》，提交铬矿储量125.5万吨。此后，1964～1985年不少单位对该矿床进行了不同程度的普查勘探和科学研究，大大加深了对该区铬矿成矿作用的认识，但找矿方面没有新的突破。
1958年，内蒙古锡林格勒盟成立582厂，开始建厂生产。至1962年，共采矿石2万吨，地表矿体已采尽。1971年，由国家投资，计划开采深部矿体，后因一些技术和经济问题，于1986年停止施工。1988年，重新投产，以采富矿石为主，1995年另外开井，采集深部富矿，年产能力3万吨，市区建有铬铁合金厂，目前一期单台炉投产，日产60＃铬铁合金25 t。销往河南、山西、辽宁等地，后精矿粉因价格问题亏损而停产。
（二）、区域地质概况
二连浩特-贺根山-黑河蛇绿岩带，大地构造位置处于西伯利亚克拉通与华北克拉通之间的天山-兴蒙造山系（任纪舜等，1999），即徐志刚等（2008）在《中国成矿区带划分方案》中划分的东乌珠穆沁旗-博克图Fe-Mo-Sn-W-Cu-Pb-Zn-Ag-Au-CrⅢ级成矿带（Ⅲ-48）（Ⅴm-Ⅰ）。该蛇绿岩带是西伯利亚板块与锡林浩特-松嫩微陆块（属广义的中朝板块）碰撞的标志，碰撞作用始于晚泥盆世，延续至早石炭世，并由此造成沿该线两侧早石炭世维宪期之前生物群有明显差异（郭胜哲，1986），在晚泥盆世出现最早的生物混生（唐克东和张允平，1991）。并由此形成了贺根山蛇绿岩带及蛇绿岩型铬矿（图3-37）。
图3-37 内蒙古中东部蛇绿岩分布示意图
1—第四系；2—侏罗系；3—石炭-二叠系；4—泥盆系；5—寒武-志留系；6—中元古界；7—太古界；8—蛇绿岩带；9—花岗岩
（引自宝音乌力吉等，2009，略作修改）
贺根山蛇绿岩带是内蒙古东北部保存较完整的一个蛇绿岩带，它不仅为我们追溯和反演已消亡的古蒙古洋盆历史提供了宝贵信息，而且也是内蒙古乃至整个华北最重要的铬矿产地。该蛇绿岩带出露长约120km，宽20km左右，面积约2158km2。蛇绿岩岩石组合较为齐全：有最接近原始地幔的二辉橄榄岩，代表亏损上地幔的斜辉辉橄岩-纯橄岩；岩浆房底部的镁铁质堆积杂岩；岩浆房主体的层状-块状辉长岩；以及岩浆房的最终分异物斜长花岗岩以及辉绿岩席状岩墙群、变质基性火山岩和硅质岩及深海沉积物盖层，完全可以和国内外典型蛇绿岩剖面对比。现已查明该蛇绿岩带中有20多个超基性岩体，自西向东包括阿尤拉海、奥约庭格勒、朝格乌得尔、赫格敖拉、赫白、白音敖拉、小坝梁、崇根敖拉、乌斯尼黑等岩体，呈S形分布。除朝格乌得尔、崇根熬拉有较大面积的超基性岩出露外，其余均为隐伏于中、新生代沉积物之下的盲岩体。据物探资料推测阿尤拉海岩体埋深最大，达900m左右，其他岩体埋深不一。其中以赫格敖拉岩体最为重要。
区内最老地层是中泥盆统，出露于岩体的西南端，为一套变基性火山岩、火山碎屑岩、凝灰岩、绿片岩夹薄层燧石、结晶灰岩。岩体东端与上石炭统地层接触，主要为酸性火山碎屑岩、凝灰岩。岩体的北、西、南三面均为下白垩统的泥砾岩、砂砾岩、粉砂岩所不整合覆盖。采自堆晶岩的放射虫硅质岩，经王乃文鉴定为晚泥盆世（鲍佩声等，1999），其形成时间应早于泥盆世，而其侵位时间大约为晚石炭世—早二叠世（内蒙古地质局，1984）。
褶皱构造仅见于中泥盆统和上石炭统地层中。中泥盆统地层组成单斜构造，上石炭统地层呈轴向近南北（区外），向北仰起的向斜。主要断裂构造以正断层为主，其次为性质不明断层。断层走向北东和近东西向，分别向北西和南倾斜，倾角中等。贺根山蛇绿岩侵位受成岩前的近东西向与北北东向两组断裂带控制。
（三）赫格敖拉岩体特征
赫格敖拉基性-超基性岩体主要由堆晶杂岩和局熔（亏损）地幔岩组成。该岩体分布于阿尤拉海赫格敖拉-乌斯尼黑S形岩体群的中段，地表呈椭圆形，东西长约9km、南北宽6km，面积大约50km2（图3-38），岩体走向北东30°，倾向120°，倾角67°～78°，据华北油田钻孔资料，曾于3000多米深处见超基性岩。物探资料推算，最深厚度可达5800 米（与内蒙古地调院会议交流，2011）。岩体四周被广厚的中、新生界地层覆盖，露头极为少见。向东与赫白岩体相连。
出露地表的堆晶岩由强蚀变的斜长石、橄榄石（镁贵橄榄石）及单斜辉石（透辉石）组成；而局熔地幔岩则由约90%的变形变质斜辉辉橄岩和10%的纯橄岩及基性岩脉组成，纯橄岩透镜体成群分布，与斜辉辉橄岩呈迅变或过渡接触关系。主要造岩矿物为橄榄石（镁橄榄石）、斜方辉石（顽火辉石），各类岩石中副矿物铬尖晶石成分变化不大（鲍佩声等，1999）。
根据岩石组合划分出4个岩相带：
1）堆晶杂岩岩相带（Ⅰ带）：分布于岩体东北缘，延长数千米，最宽 570余米，面积约0.6km2，占岩体总面积的2%。底部为含长纯橄岩、纯橄岩和浅色橄长岩，三者常互相过渡，含长纯橄岩中普遍有铬矿化。上部为橄榄辉长岩，总厚约40余米。
2）斜辉辉橄岩-纯橄岩-层状基性岩岩相带（Ⅱ带）：分布于岩体东侧，呈北东向延伸，宽度变化较大1000～2200m不等，
（斜辉辉橄岩）、φ1（纯橄岩）、v（辉长岩）3种岩石类型频繁交替出现，相当于壳-幔过渡带。该带产出一些小型铬矿床，如基东矿、620矿、820矿等及若干矿化点。
3）斜辉辉橄岩岩相带（Ⅲ带）：分布于Ⅱ带西侧，主要由斜辉辉橄岩和零星纯橄岩组成。有少量铬矿化，如B265矿化点。
4）纯橄岩-斜辉辉橄岩杂岩岩相带（Ⅳ带，赫格敖拉3756）：为岩体最主要岩相带，分布于岩体近中部，走向北东，主要由斜辉辉橄岩及拉长（具塑性变形）的纯橄岩透镜体组成，二者交互呈条带状杂岩出现，斜辉辉橄岩中辉石含量一般＜17%，构成低辉石带；两侧斜辉辉橄岩中辉石含量一般为20%～25%左右。纯橄岩异离体较为密集，呈北东向展布。带内脉岩数量少并蚀变强。这些特征显示该地段的地幔岩具有较高的部分熔融程度及较强的塑性变形，为典型的局熔地幔残渣相岩石。该带赋存了3756等主要有工业价值的中型铬矿床（图3-38）。
岩石普遍具强烈蛇纹石化，并伴有硅化和碳酸盐化，少见矿物残晶。
据内蒙古地调院资料，斜辉辉橄岩平均岩石化学成分为SiO2：35.5%～37.8%，TiO2：0～0.3%，A12O3：0.27%～0.5%，Cr2O3：0.05%～0.17%，Fe2O3：0.71%～3.18%，FeO：1.63%～4.29%，MgO：56.85%～59.28%，CaO：0～0.13%，K2O：0～0.01%，Na2O：0～0.2%。由此表明斜辉辉橄岩镁高，低钛，钾、钠极低。其中3756地段m/f比值为10.56，基性程度最高。41地段m/f比值为9.4，基性程度低。基东地段m/f比值为10.08，介于上述两者之间。斜辉辉橄岩MgO/SiO2比值与辉石含量呈反比，与橄榄石含量成正比。
图3-38 贺根山岩体地质图及剖面图
（据内蒙古地质矿产局109地质队，1980，改编）
1—第四系；2—下白垩统泥砾岩-粉砂岩；3—上石炭统火山碎屑岩-凝灰岩；4—下—中泥盆统变基性火山岩-凝灰岩夹燧石结晶灰岩；5—流纹斑岩；6—花岗斑岩；7—纯橄岩；8—斜辉辉橄岩（低辉方辉辉橄岩）；9—含长纯橄岩；10—橄长岩；11—辉长岩；12—岩石莫霍面；13—岩相带界限及编号；14—实测-推测地质界线；15—实测不整合地质界线；16—接触界线或地质产状；17—流面及产状；18—大-小型铬矿床；19—铬矿点；20—铬矿范围；21—向斜轴
纯橄岩岩石化学特征：3756地段m/f、MgO/SiO2值为9.73和2.76；41 地段这两比值最低。含矿纯橄岩与一般纯橄岩相比，Mg含量高，镁硅比值高，表明铬矿形成与富镁纯橄岩关系最为密切。
堆晶岩表现为具有轻稀土亏损和强烈的Eu正异常特点；而局熔地幔岩则具有以REE总量相对较低的Ⅴ形（或U形）（3756矿体的斜辉辉橄岩）或烟斗状（基东矿体的纯橄岩）分配形式及Eu负异常的特征。基本与我国蛇绿岩型铬矿床的岩石稀土特征相同（图3-39、贺根山各类岩石的REE分配形式）。
图3-39 岩体各类岩石的REE分配型式
（据鲍佩声等，1999）
（a）地幔岩的REE型式：1、2、3为3756区斜辉辉橄岩；4、5为3756矿区纯橄岩；6为基东矿区斜辉辉橄岩；7为基东矿区纯橄岩；
（b）堆晶岩的REE型式：8为含长纯橄岩；9为橄长岩；10为层纹状辉长岩；11为基东矿区橄长岩；12为3756矿区辉长岩
（四）矿床特征
赫格敖拉铬矿床有两种类型铬矿：一是产于蛇绿岩套剖面之岩石莫霍面下方，局熔程度较高的地幔橄榄岩（斜辉辉橄岩-透镜状纯橄岩）中的局熔型富铝铬矿床；二是产于超镁铁质-镁铁质堆积杂岩底部的岩浆堆晶型铬矿化（点）。前者以3756中型铬矿床为代表，此外，还有基东、620、820等小型矿床，具有较大的经济价值。后者以D2三角点矿化为代表，工业意义不大。现以3756 铬矿床为例，简述其矿床特征。
赫格敖拉3756铬矿床：地表长830m，宽110～300m，面积约0.13km2，向北东东方向侧伏至垂深440m（4线以北），沿轴向最大延伸已控制930m，沿倾斜最大延深280m。主要矿体群在西南部（25～27线），总体上形成一个北东走向的含矿带。由186个矿体群组成，地表出露仅有20余个，其余全部为盲矿体（图3-40、图3-41）。矿体几乎全部产于Ⅳ岩相带的透镜状纯橄岩中。矿群总体轴向50°～60°，倾向南东，倾角20°～70°。矿体形态较复杂，以透镜体居多，似脉状矿体较稳定。
矿石自然类型，以稠密浸染状为主，占62.04%；中等浸染状矿石占29.69%；致密块状矿石占4.1%；稀疏浸染状矿石占3.67%。矿体呈透镜状、豆荚状，似脉状和囊状等产出。矿体有急剧变薄，尖灭再现的现象，而且产状陡缓变化大，从而造成复杂的矿体形态。
矿体主要以纯橄岩为直接围岩，少数为斜辉辉橄岩。矿体与围岩接触关系呈清晰的迅速渐变或突变关系。近矿围岩蚀变强烈，通常为致密隐晶质-微晶质的绿泥石集合体，多数铬矿体四周发育有几厘米至几十厘米的蛇纹石-绿泥石薄壳，与纯橄岩呈渐变过渡关系。
图3-40 赫格敖拉3756铬矿床地质图
（据张国维等，1996）
Ⅴ—辉长岩及蚀变辉长岩；Mg—菱镁岩；φ2—斜辉辉橄岩；φ1—纯橄岩；1—剖面线；2—铬矿体及编号；3—铬矿点；4—产状；5—实测断层；6—推测断层
图3-41 赫格敖拉3756铬矿床B线纵剖面图
（据内蒙古地矿局109地质队，1976，简缩）
1—第四系；2—古近系—新近系上新统红粘土；3—纯橄岩；4—斜辉辉橄岩；5—斜辉橄榄岩；6—辉长岩；7—铬矿体；8—勘探线编号
铬尖晶石多呈半自形晶、他形粒状集合体散布于蛇纹石、绿泥石等脉石矿物中。粒度多大于1mm。尘点状磁铁矿常呈微粒集合体，沿铬尖晶石边缘、裂隙或与脉石矿物一起交代铬尖晶石，形成残余结构和网状结构。矿石构造包括豆斑状或瘤状构造，大部分属稠密浸染状矿石，少部分为块状矿石；浸染斑点状构造，稠密浸染状及细粒-中粒浸染状矿石具此类构造，为矿床的主要构造类型；条带状构造，条带宽2～5mm，与纯橄岩相间排列，条带为密集程度不同的铬尖晶石集合体；含块状矿石的浸染状构造，多出现在深部矿体中。
矿石矿物：金属矿物以铬尖晶石为主，尘点状磁铁矿为次，极少量黄铁矿、黄铜矿和赤铁矿。脉石矿物以叶蛇纹石为主，绿泥石次之。
铬尖晶石主要为镁质铝铬铁矿，次为富铬尖晶石和富铁铬铁矿。在不同矿石类型中，铬尖晶石化学成分变化规律为：由稀疏浸染状→致密块状矿石，铝镁及Mg/（FeO）、A12O3比值递增。矿石平均化学成分：Cr2O3：23.62%、Fe2O3：1.4%、A12O3：13.27%、SiO2：16.52%、FeO：9.48%、MgO：25.92%、CaO：0.57%、K2O：0.011%、Na2O：0.076%。
成矿物化条件：据白文吉（1986）研究，在海底环境的低压条件下，大约800℃形成堆晶岩中的铬矿体，铬矿形成的氧逸度（ƒo2）为10-6。
（五）矿床成因与成矿模式
1.成矿期次及成矿时代
贺根山铬矿床存在两种成因亚型铬矿：一种是产于蛇绿岩剖面岩石莫霍面下方，局熔-改造型富铝铬矿床（简称“局熔型”）。另一种是产于岩石莫霍面上方，超镁铁质-镁铁质堆积杂岩底部的岩浆结晶-分凝型铬矿化（点）（简称“堆晶型”）。
成矿时代：目前尚未见到铬尖晶石成矿年龄的报道，一般借用蛇绿岩套中基性-超基性岩石测年数据，代表铬矿形成时代，对于“堆晶型”铬矿而言，成岩-成矿时间比较接近。而“局熔型”铬矿，矿石的形成可能会较围岩（母岩）更早。本区采自矿区的放射虫硅质岩，经王乃文鉴定为晚泥盆世（鲍佩声等，1996）。包志伟、陈森煌等在贺根山用超基性岩全岩，获得Sm-Nd 等时线年龄403±27Ma、在乌斯尼黑斜辉橄榄岩全岩的K-Ar法年龄为430Ma和285Ma（包志伟，1994）。苗来成等（2005）在贺根山取得3个锆石同位素年龄，分别为：辉长岩295Ma、基性脉岩298ma、基性熔岩293ma，相当于晚石炭世。因此岩带的形成时间应定为晚泥盆世—晚石炭世早期较适宜。
2.矿床成因及成矿要素
贺根山铬矿床两类铬矿的成矿物质，均来源于地幔深部局部熔融的尖晶石二辉橄榄岩。其形成过程可概括如下：①当上地幔熔融程度高时，可形成镁质基性-超基性岩浆（例如玄武岩浆-苦橄岩浆），沿合适通道进入地壳，汇聚成“壳下岩浆房”分异冷凝，岩浆中晶出的铬尖晶石，因重力下沉到岩浆房的底部，在温度不迅速下降的条件下，随着晶体堆积的增多，最终可以形成具有岩浆矿床特征的堆积层状铬矿；②残留在地幔中的物质则在“对流”运动中，通过部分熔融作用，经历了逐区“分馏-精练-萃取”过程，被除去低熔点组分，从而形成以斜辉橄榄岩-斜辉辉橄岩-纯橄岩组合为代表的“地幔残渣”（亏损地幔岩）。这个过程，可视为是一种借助于“热扰动”或“动力扰动”形式出现的壳-幔分异作用。在此过程中，地幔岩中的含铬辉石及副矿物铬尖晶石发生相变，形成富铬矿浆，矿浆在熔体中最初为分散的小熔滴，受地幔塑性流动和重力作用运移并逐渐变大，从而进一步聚集形成铬矿，即“局熔型”铬矿矿石。鲍佩声等（1999）通过对矿物变形研究，贺根山岩体古应力差和应变速率均较低。对铬矿浆熔滴的汇聚是不利的。
（六）找矿模型及预测要素
1）寻找镁质超基性岩是最重要的找矿前提。断裂活动对岩体的侵位有重要的作用，它既是岩体侵位的基本原因，也可能破坏原始壳-幔过渡带的结构。应注意那些可能标志早期洋壳主要遗迹的消减带或缝合带构造带，其次是经过次级扩张又有褶缩关闭的弧后盆地褶皱地带，特别是其边缘的一些断裂带，往往有利于成矿；
2）通过对蛇绿岩套火山岩性质的研究，对“残片”所代表的原始洋底构造部位和性质做出分析和判断，这将有利于寻找曾有过高热流活动构造部位的“遗迹”——成矿有利地区；
3）岩体侵位是一种随机截取，因此，岩体的形态、规模、产状都不是决定因素。岩体的含矿性，取决于早期“壳-幔分异”作用的程度和性质。分异好的，具有以下两个特点：一是保存良好分异的“壳下岩浆房”（堆晶岩发育），二是岩相分带清楚。在查清岩相分带特征的基础上，尽可能恢复原始壳-幔过渡带，进而寻找新的铬矿体。
4）鉴于3756矿床大部分为盲矿体，经实践证明磁法圈定岩体边界效果较好，重力、扭秤仅在几个已知铬矿体上有反映，而电法、化探效果不佳。如5号矿体近南北向重、磁异常综合剖面（图342）所示，重力、扭秤均有明显的异常反映，∆g=0.16×10-5m/s2，Ⅴxz=60×10-91/s2，Ⅴ∆=116×10-91/s2。∆g曲线基本对称，矿体上显示重力高、磁异常由低到高的过渡带。
图3-42 3756铬矿体上重、磁异常综合剖面图
（据廖昌庆等，1996）
Q—第四系；φδ—纯橄岩；φω—斜辉辉橄岩
如图3-43 所示，620矿体走向北东30°，由3个较大的透镜状铬矿体组成，走向40m，延伸25m，倾向东南，倾角较陡。矿顶埋深4～5m，矿石以细粒稠密浸染状为主。在矿体上反映为重力异常，而磁异常特征不明显。然而在地形条件较复杂的地区，以上方法会受到一定限制。因此，有必要进一步探索新方法，以期取得寻找盲矿体的更好效果。
二、内蒙古乌拉特中旗索伦山铬矿床
索伦山式铬矿床产于西伯利亚古板块与中朝古板块的缝合带西段——索伦山蛇绿岩带。其形成时间应早于泥盆世，而其侵位时间大约为晚石炭世—早二叠世。岩体由亏损地幔橄榄岩和堆晶岩组成。铬矿毫无例外地赋存于各类岩相的纯橄岩中。其中，地幔岩局熔改造型豆荚状铬矿具工业价值，矿石类型因不同岩体而异，总体而论仍为高铝型铬矿，矿床规模属小型。
（一）概况
索伦山铬矿，产于内蒙古巴彦淖尔市乌拉特中旗北东90km处，地理坐标：东经117°18＇00〞，北纬49°26＇00〞。该地区产出索伦山（察汗胡勒、阿布铬、乌珠尔）等铬矿床（点）。有工业价值的矿体主要分布于索伦山岩体西段的察汗胡勒矿区、中段的察汗奴鲁矿区及东段的土格木矿区。多数为耐火级铬矿矿石，少为冶金级铬矿石。累计探明C级储量10.0万吨，D级39.2万吨，共计49.2万吨；截止1993年，保有C级储量7.7万吨，D级38.3万吨，共计46.0万吨，为小型铬矿床。
图3-43 620矿体重磁综合剖面图
（引自廖昌庆等，1996）
TCr—倒塌的铬矿；1—第四系；2—斜辉辉橄岩；3—辉长岩；4—断层；5—钻孔；6—浅井
该矿床于1957年由内蒙古地质局区调队发现，1958年航磁及地幔磁测圈定超基性岩体。1958～1963年内蒙古地质局组织普查勘探，提交储量报告。期间，北京、天津部分科研单位进行了相关科学研究工作。1962年开始零星开采，1982年由乌特拉中旗计委投资建厂，至1994年采出矿石12万吨，精矿粉送天津、河北化工厂，部分富矿销往太钢。
（二）区域地质概况
索伦山蛇绿岩带西起哈布特盖，向东经索伦敖包、阿不盖敖包（即阿布格）、乌珠尔到哈尔陶勒盖等，为一个东西长约100多千米的狭长地带，已查明岩带中有5个较大的超基性岩体：索伦山、阿布格、乌珠儿、平顶山、哈也（图3-44），分布于索伦敖包-满都拉大断裂北侧，与围岩均为断层接触。多数学者认为该岩带东延可与西拉木伦蛇绿岩带连接（李锦轶等，1987，1998）；陆松年等指出，兴蒙造山系中的二连-贺根山洋盆到早泥盆世已闭合，额尔古纳、松辽、佳木斯等地块已拼贴联合，此时，古亚洲洋东段仅保存西拉木伦洋盆。到石炭纪中期至早二叠世，洋壳消减和大陆增生已近尾声，最重要的地质事件是西伯利亚与华北全面碰撞对接，索伦山-西拉木伦是西伯利亚古板块与中朝古板块的分界线，也是古亚洲洋最终消亡的地带。一系列新的研究成果表明，古亚洲洋最终封闭的时间为二叠纪末（248±4ma）（陆松年，2010，内部资料）。据《中国成矿区带划分方案》（徐志刚等，2008），索伦山铬矿床地处白乃庙-锡林浩特Fe-Cu-Mo-Pb-Zn-Mn-CrⅢ级成矿带（Ⅲ-49）（Ⅴm-Ⅰ）。
蛇绿岩带形成时代为志留—泥盆纪，据陶继雄等（2004）取岩体中橄榄辉长岩，测得单颗粒锆石U-Pb年龄值为（433.6±3.6）Ma，侵位于石炭—二叠纪。
（三）岩体特征
索伦山、阿布格（即阿布盖敖包）和乌珠儿岩体是索伦山蛇绿岩带中较大且矿化较好的3个构造岩块。
索伦山岩体东西长32km，南北宽2～6km，面积72km2，向北进入蒙古境内，由斜辉辉橄岩和纯橄岩组成，局部含少量二辉橄榄岩。纯橄岩以复杂多变的透镜状、条带状、似脉状体广泛分布于斜辉辉橄岩中，含量约占10%，主体为斜辉辉橄岩，二者界线截然。岩体岩相分带不明显，仅在索伦山主峰以北见少量斜辉辉橄岩-二辉橄榄岩，其他均为斜辉辉橄岩-纯橄岩岩相。
阿布格岩体位于索伦山岩体东28km，岩体东西长9.5km，宽0.18～3.5km，主要由纯橄岩、斜辉辉橄岩组成，含少量异剥橄榄岩、辉石岩、辉长岩等透镜体。纯橄岩出露占岩体总面积30%，从南到北，以纯橄岩-斜辉辉橄岩岩相带、纯橄岩岩相带交替出现。岩体地表风化甚强，主要为硅化和碳酸盐化，蚀变深度可达数十米。
乌珠儿岩体位于阿布格岩体东18km，东西长3.5km，宽0.7～1.5km，岩体由纯橄岩-斜辉辉橄岩、辉长岩及枕状熔岩组成，为该区出露较完整的蛇绿岩组合，但缺少岩墙群。可分为纯橄岩岩相带、纯橄岩-斜辉辉橄岩杂岩岩相带及二辉橄榄岩3个岩相带。
3个岩体岩石化学成分差别不大，相比较而言，索伦山岩体铬含量较阿布格和乌珠儿略高。岩体蛇纹石化均较强，特别是阿布格岩体，除蛇纹石化外还具有强烈的硅化和碳酸盐化。
（四）矿床地质特征
索伦山岩体中绝大部分工业矿体，无例外地产于斜辉辉橄岩-纯橄岩杂岩带之纯橄岩透镜体中，并有成群出现、成带分布特点，且以北部居多，往南减少。自西向东包括察汗胡勒、察汗奴鲁和土克木3个矿区（图3-44，图3-45），共发现100多个矿体，其中近80%为盲矿体。矿体形态复杂多样，以透镜状、豆荚状、脉状、网脉状、囊状、巢状、筒状及不规则状等为主。规模可由数十至300m。矿石类型以浸染状为主，造矿铬尖晶石为铬矿（富铬型），Cr2O3变化较大，可由10.68%～31.59%，个别可大于 40%～50%；Cr2O3/＜FeO＞可达 3.03～3.06。探明储量占全区（索伦山全岩体）73.45%。
图3-44 索伦山岩体察汗胡勒矿区地质图
（据内蒙古地矿局地质研究队，1984）
1—斜辉辉橄岩；2—纯橄岩；3—铬矿体；4—断层；5—矿体产状；6—勘探剖面线
阿布格岩体矿化较为普遍，主要有两个矿群Cr209、Cr207，共发现37个矿体，地表仅出露4个，矿体一般长数十至百余米，厚数米。探明储量占全区9.8%。
乌珠儿岩体有Ⅰ、Ⅲ两个矿群，由18个矿体组成，仅有2个矿体出露地表，其他均为盲矿体。探明储量占全区16.9%。
矿石矿物成分主要为铝铬铁矿，较之贺根山铬矿含铁略高，可能与蚀变作用较强有关。此外，含少量磁铁矿、赤铁矿、磁黄铁矿和镍黄铁矿。脉石矿物以叶蛇纹石为主，绿泥石、菱镁矿次之。
矿石呈他形-半自形晶结构、碎裂结构、碎斑结构、交代结构、网脉状结构及塑性变形结构。矿石构造复杂，主要为浸染状矿石，其次为致密块状、条带状、巢状、囊状、斑杂状、反斑杂状构造。
图3-45 索伦山察汗奴鲁矿区Ⅱ矿群4勘探线剖面图
（据内蒙古地矿局地质研究队，1984）
1—斜辉辉橄岩；2—纯橄岩；3—碳酸盐化超基性岩；4—辉绿岩；5—铬矿体；6—推测断层；7—岩相界线；8—钻孔
（五）找矿建议
内蒙古地区是我国较早开展铬矿找矿工作地区之一，并在贺根山—索伦山一带找到了中型矿床一处；小型矿床2处；矿点、矿化点36处，其中近80%为盲矿体。足以证明该地段具有良好的找矿潜力。提出以下几点找矿建议：
1）内蒙古具工业价值的铬矿均为蛇绿岩型铬矿。索伦山蛇绿岩带又是西伯利亚古板块与中朝古板块的缝合带。因此，应加强区域蛇绿岩带的研究，寻找成矿有利地段，加大找矿勘查力度，有希望取得新进展。
2）加强方法研究。找矿方法的成功，往往可以带动找矿突破。鉴于内蒙古第四系覆盖面积较大，以往找到的矿体，大部分为盲矿体，积累了一定的经验，很有必要进一步加强研究物探综合方法在寻找铬矿方面的应用研究，力争有所突破。
3）关注国内外新的成矿理论，深入研究中国铬矿成矿规律，并结合深部找矿的趋势，重点研究深部找矿的技术方法，以点带面。
『伍』 矿产资源/储量边界线的种类
矿产资源/储量边界线因其含意不同而名称各异，矿产资源/储量边界的圈定一般先在单个工程内确定矿产资源/储量边界 ( 或边界基点) ，然后再根据所有工程内的资源/储量边界基点，在剖面上或平面上连结对应的点形成界线而得到矿体在三维空间的各种边界。
1. 零点边界线
零点边界线是在投影面上，矿体厚度或有用组分含量趋于零的各点连线。即矿体尖灭点的连线。零点边界线常常是为了确定可采边界线时的辅助线，而不是真正意义的资源/储量边界，因矿产储量不可能计算到零点边界上。
2. 可采边界线
可采边界线是按最小可采厚度和最低工业品位或最低工业米百分值所确定的基点的连线，它是用来圈定工业矿体的边界位置，即可采边界内的矿量为储量或基础储量。
3. 矿石品级和类型边界线
在可采边界线的范围内，按矿石技术品级和类型的要求标准，划分的不同技术品级和矿石类型的分界线。表明各种品级和类型的矿石在工业矿体中的分布情况。
确定矿石品级和自然类型的边界时，是在可采边界范围内，必须注意控制矿石品级和自然类型的地质因素。例如在确定氧化带与原生带的边界时，必须考虑氧化带和原生带的界线主要由地下水位控制，而地下水面在较短的距离内可以视为水平的，因此其边界应是水平的。
表 8-7 一般有色金属矿石自然类型划分标准表
铁矿石原生矿与氧化矿的划分，一般是按 TFe/FeO 的不同比值来衡量。当铁矿床中含铁矿物主要是磁铁矿，后经氧化成赤铁矿、褐铁矿时:
原生矿石: TFe/FeO ＜ 2. 7; 混合矿石: TFe/FeO2. 7 ～ 3. 5; 氧化矿石: TFe/FeO＞ 3. 5。
当矿石中含铁矿物主要是菱铁矿或硅酸铁比较高的磁铁矿矿石，原生矿、氧化矿的划分标准另行考虑。
4. 储量类别边界线
即按不同储量类别条件所圈定的界线，例如储量、基础储量和资源量的分界线。
5. 内边界线与外边界线
内边界线是矿体边缘见矿工程控制点连接的界线，它表示被勘查工程所控制的那部分矿体的分布范围;外边界线是根据边缘见矿工程向外或向深部推断确定的边界线，以表示矿体的可能分布范围。从空间上说，零点边界线属于外边界线，而其他几种边界线可在内边界线之内，也可在内、外边界线之间。
6. 暂不能开采边界线
这条界线是根据边界品位圈定的，此线与可采边界线之间的矿量为资源量。
『陆』 火成岩的物质成分
火成岩的物质成分是火成岩的最基本特征，它既是火成岩分类命名的基本依据，也是研究岩浆起源、生成和演化的重要手段。
一、火成岩的化学成分
研究表明，地壳中的所有元素在火成岩中都有出现，但含量差别很大。根据元素在火成岩中的含量和地球化学意义，将其划分为主要元素、微量元素和同位素。
(一)主要元素
组成火成岩的元素有很多种，但以O、Si、Al、Fe、Mn、Mg、Ca、Na、K、Ti、P、H等12种元素为主，并以O元素含量最高，达45%以上。这12种元素的总和占火成岩总质量的99%以上，称为主要造岩元素。在研究火成岩时，不是以元素的形式表示其化学成分，而是以氧化物的形式表示，即SiO2、TiO2、Al2O3、Fe2O3、FeO、MnO、MgO、CaO、Na2O、K2O、P2O5、H2O。这些氧化物在火成岩中的含量通常大于0.1%，称为主要造岩氧化物(表2－1)。根据研究需要，还可以给出CO2、Cr2O3等含量。
表2－1 中国火成岩的化学成分(wB/%)
(1)SiO2是火成岩中含量最多、变化范围最大，也是最重要的氧化物。根据SiO2含量将火成岩划分成酸性岩(SiO2＞63%)、中性岩(SiO252%～63%)、基性岩(SiO245%～52%)和超基性岩(SiO2＜45%)四种类型。通常以SiO2含量高低来称谓火成岩的酸性或基性程度，含量越高者，岩石酸性程度就越大，基性程度就越低。对火成岩化学成分研究表明，随SiO2含量的变化，其他主要造岩氧化物含量发生规律性变化(图2－1)。随SiO2含量增加，Na2O、K2O含量逐渐增加，FeO、MgO含量不断减少;而CaO、Al2O3含量由超基性岩到基性岩随SiO2含量增加而快速增加，之后由基性岩向中性岩、酸性岩变化时则逐渐减少。岩浆中的SiO2在组成长石、云母、角闪石、辉石等硅酸盐矿物之外，还有多余时，就以独立的石英矿物出现。
图2－1 火成岩中SiO2含量与其他氧化物之间的关系(邱家骧，1985)
(2)Na2O与K2O含量之和称为全碱含量，它在不同的岩性中有较大差别(表2－1)。Na2O和K2O是碱性长石的主要成分，全碱含量较高时，岩石中可以出现碱性暗色矿物和副长石。在火成岩研究中，常用里特曼指数(σ)划分岩石的碱性程度，σ=［w(Na2O+K2O)2］/［w(SiO2)－43%］，σ＜3.3的岩石为钙碱性岩，σ=3.3～9的岩石为碱性岩，σ＞9的岩石为过碱性岩。但对于SiO2含量很高的岩石(SiO2＞70%)，里特曼指数在确定碱性还是亚碱性时显得无效(邓晋福等，2004)，这是因为SiO2的稀释效应会导致碱含量相对偏低，计算出的σ值偏小，会被误定为钙碱性岩系。如某些SiO2＞80%的碱性流纹岩，应特别注意。
(3)Al2O3是仅次于SiO2的造岩氧化物，火成岩中Al2O3含量主要在10%～18%之间。Al2O3与SiO2及CaO、Na2O、K2O结合形成斜长石、碱性长石和似长石等矿物;与FeO、MgO、CaO和SiO2结合形成辉石、角闪石和黑云母等矿物。Al2O3同样在火成岩分类和成因研究中具有重要作用:①根据碱含量同CaO和Al2O3含量之间的相对比值，将火成岩划分为过碱质岩石(Al2O3＜Na2O+K2O，分子数，下同)、过铝质岩石(Al2O3＞CaO+Na2O+K2O)和偏铝质岩石(Na2O+K2O＜Al2O3＜CaO+Na2O+K2O);②在亚碱性系列玄武岩中，将Al2O3≥16%(质量分数)的岩石称为高铝玄武岩;③将铝指数A/CNK=Al2O3/(CaO+Na2O+K2O)(分子数比)＞1.1的花岗岩，称为S型花岗岩。
(4)MgO、FeO与SiO2结合形成铁镁硅酸盐矿物，如橄榄石、辉石等。因MgO、FeO与SiO2含量呈负相关(图2－1)，因而只有在SiO2含量低的情况下，才出现橄榄石、辉石。依据主要元素进行火成岩岩石系列划分、岩石分类和成因研究，是火成岩研究的主要方法之一，所涉及的内容较多。对于初学者来说，以下三个有关主量元素的应用应该掌握和了解。
1.火成岩岩石系列的划分
火成岩可以分成三个岩石系列，即碱性、钙碱性和拉斑玄武岩系列，后两者合在一起称为亚碱性系列。首先根据硅碱图(图2－2)，区分碱性系列(A)和亚碱性系列(S)。对于亚碱性系列岩石，利用TFeO/MgO－SiO2图解及TFeO/MgO－TFeO图解(图2－3)或者AFM图解(图2－4)进一步区分是拉斑玄武岩系列，还是钙碱性系列。也可依据SiO2－K2O图解把亚碱性系列的火成岩区分为低钾、中钾、高钾和钾玄岩类型(图2－5)。
图2－2 硅碱图解(Irvine，1977)
图2－3 火山岩系列划分的TFeO/MgO－SiO2图解(a)、TFeO/MgO－TFeO图解(b)(Miyashiro，1974)
图2－4 火成岩系列划分的AFM图解(Rollison，1993;转引自杨学明等，2000)
图2－5 亚碱性火成岩系列划分的SiO2－K2O图解(LeMaitreetal.，1989;Rickwood，1989)
有关火成岩系列划分还有许多其他图解，应用时要特别注意每个图解的使用条件，不能生搬硬套。例如:应用硅－碱图解划分岩石系列时，对于高硅的花岗岩和流纹岩(一般SiO2＞70%)需小心，因SiO2含量高导致碱含量低，与里特曼指数一样，在确定是碱性还是亚碱性系列时会无效(邓晋福等，2004)，使得碱性花岗岩和碱性流纹岩落入亚碱性系列区，这显然是错误的。原图解中两个系列分界线的上端点终止于SiO2含量为67%处，没有向上延伸也正是此原因。Irvine＆Baragar(1971)给出了该图解分界线的数学方程式为:S=－(3.3539×10－4)A6+(1.2030×10－2)A5－0.15188A4+0.86096A3－2.1111A2+3.9492A+39。式中S=w(SiO2)、A=w(Na2O+K2O)，当岩石中SiO2大于由公式算出的S时为亚碱区，反之为碱性区。邓晋福等(2004)建议对于SiO2＞70%的火成岩系列划分时采用Wright(1969)提出的碱度率［AR=w(Al2O3+CaO+Na2O+K2O)/w(Al2O3+CaO－Na2O－K2O)］，并用SiO2－AR图解加以区分(图2－6)。同时，岩石中出现碱性暗色矿物是最重要的岩相学标志，过碱指数(［(Na2O+K2O)/Al2O3］＞1，分子数)是鉴别碱性花岗岩(流纹岩)的最可靠地球化学参数。
图2－6 火成岩划分法的SiO2－AR关系图解(Wright，1969;转引自邓晋福等，2004)
研究表明，不同系列的火成岩，其岩浆起源、演化和形成的构造背景存在许多差异，因此，准确的岩石系列划分有助于火成岩成因的厘定，相关内容将在后续的章节中介绍。
2.Harker型岩石化学成分变异图解
这是最简单但又被经常使用的一种图解，该图常以SiO2或MgO含量为横坐标，其他主要氧化物含量为纵坐标构成的图解(图2－7)。根据研究需要，也可以选择相关参数，如碱度率(AR)、分异指数(DI=Q+Or+Ab+Ne+Lc+Kp，标准矿物)等作为变量加以研究。一般所使用的氧化物数据应该是将硅酸盐全分析中的H2O、烧失量等去除后重新换算出的氧化物含量。这种图解表示出随SiO2或MgO含量的变化，其他氧化物或参数的变化趋势(图2－7)。通常，在同一个地区、空间上密切共生、成分变化较大的火成岩，如果其化学成分在Harker图解上存在较强的线性相关性，表明这些岩石很可能是同源岩浆演化形成的一组岩石。若不具相关性，意味着它们可能是不同岩浆结晶的产物。
3.CIPW标准矿物计算及应用
当火成岩在快速冷凝条件下形成时，其结晶矿物颗粒细小，或部分甚至是全部由玻璃质组成(如许多火山岩)，那么该岩石的实际矿物成分及含量就无法知晓，依靠实际矿物成分及含量的岩石分类定名就无效。为解决这一问题，人们提出了利用化学成分计算火成岩中的理想矿物组成及含量的方法，即标准矿物计算方法。目前，得以广泛应用的计算方法是由美国的Cross，Iddings，Pirsson，Washington(1902)共同提出的方法，简称CIPW标准矿物计算法。
这种方法是以无水岩浆中矿物结晶顺序的实验研究成果为依据，依次按理想分子式配成标准矿物。首先将岩石的氧化物质量百分数换算为氧化物分子数，然后按照一定的顺序，再将分子数依据一定的规律，组合成若干种理想成分的标准矿物分子，最后将标准矿物分子数换算为标准矿物质量百分数。详细的计算流程参见邱家骧(1985)主编的《岩浆岩岩石学》和林景仟(1995)主编的《火成岩岩类学与岩理学》。现在，已经编制出了相关软件，通过计算机得以快速完成。CIPW标准矿物能够概略地反映出岩石的矿物组成，但并不一定是岩石中实际出现的矿物。其计算结果被用于岩石分类(图2－8，图2－9)、岩浆形成或结晶温压条件的确定(图2－10)等诸多方面。
图2－7 美国俄勒冈州Mazama山火山岩Harker图解(Winter，2001)
图2－8 玄武岩的标准矿物分类图解(邱家骧，1988)
图2－9 花岗质岩石的标准矿物An－Ab－Or分类图(Rollison，1993;转引自杨学明等，2000)
图2－10 酸性岩浆来源深度的确定(Winter，2001)
(二)微量元素
微量元素是指那些在岩石中含量甚微的元素，其含量只能以百万分之几(10－6)表示，一般情况下，它们的总量＜1%。微量元素研究已成为现代岩石学的一个关键组成部分，比主要元素能更有效地区分岩石成因演化过程。经常提到的痕量元素有钒(V)、钴(Co)、镍(Ni)、铬(Cr)、铷(Rb)、锶(Sr)、钡(Ba)、铯(Cs)、钍(Th)、铀(U)、锆(Zr)、铪(Hf)、铌(Nb)、钽(Ta)和稀土元素(REE)等。微量元素通常不以独立的矿物出现，主要是以类质同象形式替代矿物中的主要元素，如Cr、Ni可替代橄榄石和辉石中的Mg、Fe位置，Sr可占据斜长石中Ca的位置等;其次是存在于快速冷凝的火山玻璃和气液包裹体中;第三是吸附在矿物表面。
火成岩的微量元素常常随主要造岩元素含量的变化而有规律的变化。例如，随岩石酸度的增高，亲铁元素(V、Cr、Co、Ni等)的含量降低，而碱金属微量元素(Li、Rb、Cs)随之增高。对微量元素特征的研究，可以获得有关岩石系列划分、成因和演化方面的重要信息。
稀土元素包括原子序数为57～71的镧系元素:镧(La)、铈(Ce)、镨(Pr)、钕(Nd)、钷(Pm)、钐(Sm)、铕(Eu)、钆(Gd)、铽(Tb)、镝(Dy)、钬(Ho)、饵(Er)、铥(Tm)、镱(Yb)、镥(Lu)，另外，通常也将原子序数为39的钇元素(Y)算作稀土元素。除Pm为人工放射性产物外，其余都是地球化学性质相近、难熔而共生、于次生作用中不易发生迁移的元素。稀土元素总量、曲线配分模式和铕(Eu)异常等，都蕴含着岩浆起源和演化、岩石形成机理等方面的重要信息。
在岩浆结晶作用过程中，有些微量元素优先进入结晶矿物相中，或当源区岩石发生部分熔融形成岩浆时，它们易于残留在源岩矿物中，这些元素称为相容元素;相反，在岩浆结晶作用过程中，不被早晶出的矿物捕获或容纳，而富集于残余熔体中，或当源区岩石发生部分熔融形成岩浆时，它们优先进入熔体相中，这些元素称为不相容元素，也叫湿亲岩浆元素。值得注意的是，元素相容性和不相容性的程度在不同岩浆或矿物中是有差别的。例如，P在地幔岩浆中是不相容元素，但在地壳花岗岩浆中，即便是以微量元素的形式出现，也是相容元素;再比如Cr、Ni、Co元素对橄榄石而言是相容元素，而对斜长石而言，则属于不相容元素。
不相容元素依据场强(电荷/半径比值，即离子势)大小，进一步划分高场强元素(HFSE)和低场强元素(LFSE)。离子势大于2.0的元素称为高场强元素，包括镧系元素、Sc、Th、U、Pb、Zr、Hf、Ti、Nb、Ta等;离子势小于2.0的元素称为低场强元素，也称大离子亲石元素(LILE)，包括Cs、Rb、K、Ba、Sr，二价的Eu、Pb等。
正是由于微量元素的上述岩石地球化学差异，导致了其在岩石圈纵向剖面上出现了强烈的分异现象。例如，通过岩浆作用形成的地壳，其不相容元素的丰度远高于地幔。流体对地幔的局部交代作用可以引起地幔不相容元素的富集，造成地幔成分的不均一性。来源于不同源区的岩浆，必然会保留源区微量元素的痕迹。因此，通过对火成岩微量元素特征的研究，可以揭示岩浆源区性质、岩浆演化等岩石成因信息。在火成岩成因研究中，经常应用微量元素比值及其图解，以及微量元素蛛网图(图2－11)、稀土元素配分模式图等进行示踪(图2－12)。图2－11是洋中脊玄武岩(亏损地幔源区)、碱性洋岛玄武岩(富集地幔源区)和岛弧钙碱玄武岩(流体交代地幔源区)的微量元素标准化蛛网图，三者区别明显。岛弧钙碱玄武岩亏损高场强不相容元素，尤其是亏损Nb、Ta;碱性洋岛玄武岩的Nb、Ta则强烈富集;洋中脊玄武岩则亏损大离子亲石元素Ba、Rb、K。说明这三种玄武岩浆源区物质组成的不同。图2－12是不同来源火成岩的稀土元素配分曲线图(徐夕生和邱检生，2010)，徐夕生和邱检生(2010)根据其他学者的研究成果，总结为:玻安岩是岛弧地区由被俯冲洋壳释放的流体交代后的亏损地幔直接熔融产生的岩浆结晶形成，未经历明显演化，稀土元素总量低，轻重稀土元素无明显分馏，重稀土元素略富集，具U形曲线特征;埃达克岩(即岛弧英安岩)是由俯冲洋壳(及其沉积物)直接熔融形成，轻重稀土元素分馏强烈，稀土元素曲线向右陡倾，重稀土元素强烈亏损;若这种洋壳熔融的熔体与地幔橄榄岩反应，可演化成高镁安山岩。图2－12中的“常见弧安山岩”是由玄武岩浆结晶分异形成的，具有明显的负Eu异常。有关微量元素的详细阐述内容请参阅《火成岩微量元素岩石学》教材(李昌年，1992)和《岩石地球化学》一书(Rollison著，1993;杨学明等译，2000)。
图2－11 不同构造环境中玄武岩的微量元素蛛网图(Blattetal.，2006)
图2－12 不同成因类型火成岩的稀土元素配分模式图
(三)同位素
同位素在火成岩研究中已得到广泛应用，它不仅可以确定火成岩的形成时代，还可以示踪岩浆源区性质和火成岩形成演化的过程，探讨壳幔相互作用方式及大陆地壳生长等重要科学问题。同位素可以分为稳定同位素和放射性同位素两大类。
◎稳定同位素:火成岩中应用较多的是O、H和S同位素，对它们的研究可以得到火成岩成因、岩浆起源的信息，通常应用的数据有氧同位素的δ18O值、氢同位素的δD值和硫同位素的δ34S值。以氧为例，氧同位素由16O、17O、18O组成，在地质作用和岩浆作用过程中，16O和18O由于质量差别较大而发生分馏，造成岩石圈不同位置的16O和18O组成存在差异。通常以δ18O(‰)表示氧同位素的组成，δ18O=1000×［(18O/16O)样品－(18O/16O)标准］/(18O/16O)标准，(18O/16O)标准值通常采用海水平均值。虽然地幔的氧同位素组成存在较小的不均一性，但δ18O值基本为5.7‰±0.3‰左右。不同成因的火成岩其氧同位素不同，如由变质沉积岩熔融形成的花岗岩，其δ18O＞10‰;由幔源岩浆分异形成的花岗岩，其δ18O＜6‰。
◎放射性同位素:火成岩中放射性同位素有K－Ar、U－Pb、Rb－Sr、40Ar－39Ar、Sm－Nd、Re－Os、Lu－Hf等，主要用途是确定火成岩形成年龄、示踪岩石成因及地壳的形成与演化。同位素地质年代测定中最常用的年龄计算方法是等时线年龄、模式年龄、U－Pb一致线年龄和不一致线年龄，相关内容均有专著、教材介绍，可参阅陈岳龙等(2005)编著的《同位素地质年代学与地球化学》。
在同位素示踪火成岩成因方面，通常应用的数据有Sr同位素初始比值(87Sr/86Sr)i、87Sr/86Sr、143Nd/144Nd、εNd(t)、187Os/186Os、206Pb/204Pb、208Pb/206Pb、177Hf/176Hf、εHf(t)等。火成岩同位素比值之所以可以示踪源区特征，是因为常用的这些同位素对之间质量相差太小，致使这些同位素对不可能受控于晶体－液体平衡过程而发生分馏，它们在随后的分异作用过程中保持恒定。因此，部分熔融作用形成的岩浆就具有源区同位素成分的特点。这一事实引起了同位素地球化学两个方面的主要发展，一是特定的源区以其特征的同位素组成能够被识别。图2－13和图2－14给出了亏损地幔、原始地幔、富集地幔、上地壳、下地壳等不同源区的Pb、Sr、Nd同位素组成，其差别十分明显。如由幔源火成岩(如辉长岩)部分熔融形成的花岗岩，其(87Sr/86Sr)i＜0.707;由壳源泥质变质岩(如云母片岩、富铝片麻岩)部分熔融形成的花岗岩，其(87Sr/86Sr)i＞0.708。二是同位素组成各异的源区间的混合作用、混染作用能够被鉴别。例如，图2－15是苏格兰斯凯岛古近纪－新近纪火山岩的初始Pb同位素组成，该区的酸性花岗岩和基性火山岩在图解中均呈线性排列，且位于麻粒岩相下地壳和赫布里底群岛(洋岛)地幔Pb同位素组成之间，而偏离上地壳(图2－15a)。因此，该套火山熔岩被解释为来源于地幔的岩浆遭受了麻粒岩相下地壳的混染作用而形成。
图2－13 不同源区的Pb同位素组成(Rollison，1993;转引自杨学明等，2000)
图2－14 不同地幔源区Sr－Nd同位素组成(Winter，2001)
图2－15 苏格兰斯凯岛古近纪－新近纪火山岩的初始Pb同位素比值图解(Thompson，1982)
二、火成岩的矿物成分
(一)火成岩中矿物的分类
火成岩中的矿物成分既反映岩石的化学成分，又表征岩石形成的温度、压力和流体条件;既是岩石分类命名的主要根据，又是判断岩石生成条件的重要标志。在火成岩中发现的矿物种类较多，但常见的矿物只有20多种，其中最主要的、对岩石分类起重要作用的矿物有:橄榄石族、辉石族、角闪石族、云母族、碱性长石、斜长石、石英和似长石(霞石、白榴石)等，这些矿物称为主要造岩矿物。在火成岩研究过程中，人们根据矿物的化学成分、颜色、含量、成因及在分类命名中所起的作用等，对矿物进行分类。主要有以下几种分类方法:
1.矿物的成分分类
根据组成矿物的化学成分，将火成岩中的矿物分为铁镁矿物和硅铝矿物。
◎铁镁矿物:矿物中MgO、FeO含量较高，主要有橄榄石(镁橄榄石、贵橄榄石和铁橄榄石)、斜方辉石(紫苏辉石、古铜辉石、顽火辉石)、单斜辉石(普通辉石、透辉石、易变辉石和富钛辉石)、角闪石(普通角闪石为主)、黑云母等。它们在岩石中呈现黑色、黑绿色、黑褐色等深色色调，故又称暗色矿物。富含Na2O的暗色矿物称为碱性暗色矿物，如霓石、霓辉石、钠闪石和钠铁闪石。
◎硅铝矿物:不含MgO、FeO，富含SiO2、Al2O3的矿物，主要是石英、斜长石、碱性长石和似长石。它们在岩石中呈现无色、灰白色等浅色色调，因此又称浅色矿物。
暗色矿物在火成岩中的体积百分含量称为色率，是火成岩分类和鉴定的重要标志之一。色率＞90的火成岩为超镁铁质岩，基性岩的色率为40～90，中性岩的色率为15～40，酸性岩的色率＜15。
2.矿物的含量及作用分类
根据矿物在火成岩中的含量及其在岩石分类命名中作用，将火成岩中的矿物分为主要矿物、次要矿物和副矿物。
◎主要矿物:在岩石中含量高，且对岩石类型的划分起主要作用的矿物。例如，花岗岩中的石英、碱性长石、斜长石均是主要矿物;辉石和斜长石则是辉长岩的主要矿物。
◎次要矿物:在岩石中含量少于主要矿物，对岩石大类的划分不起主要作用，但对岩石种属的确定起决定作用的矿物。例如:辉长岩中可以出现少量石英，石英的出现与否并不影响辉长岩这一大类岩石的命名，但对其是否叫做石英辉长岩或含石英辉长岩起着控制作用，所以石英在辉长岩中是次要矿物。
◎副矿物:在岩石中含量通常＜1%，不影响岩石的分类命名。常见的有磁铁矿、钛铁矿、榍石、锆石、磷灰石、褐帘石、独居石等。
3.矿物的成因分类
根据火成岩中矿物的成因，将其分为原生矿物、成岩矿物和次生矿物。
◎原生矿物:岩浆在冷凝结晶过程中形成的矿物，火成岩中大多数矿物均属此类。原生矿物按生成环境可进一步分为高温矿物和低温矿物。通常来说，火山岩岩浆因温度高，所形成的矿物属高温型，如高温斜长石、高温石英(β－石英)和高温碱性长石(透长石)等;深成侵入岩中出现低温矿物，如低温斜长石、低温石英(α－石英)和低温碱性长石(正长石)。
◎成岩矿物:在岩浆结晶结束后，由于温度、压力的不断降低，使原生矿物发生转变形成新的矿物，该矿物称为成岩矿物。例如，高温的β－石英转变为低温的α－石英;高温的透长石转变为低温的正长石;正长石发生分解形成新的条纹长石;其中，α－石英、正长石和条纹长石均属成岩矿物。
◎次生矿物:属岩浆期后矿物，是岩浆成岩以后，因受残余的挥发分和岩浆期后热液流体的交代及充填作用而形成的新矿物，次生矿物主要是流体交代原生矿物和成岩矿物形成的新矿物，或充填在粒间空隙及气孔中的新矿物。①交代原生矿物和成岩矿物所形成的新矿物也称蚀变矿物，其主要是以水化和碳酸盐化为主。例如，斜长石遭受交代作用形成钠长石、方解石和黝帘石;单斜辉石蚀变成阳起石、透闪石;黑云母转变成绿泥石。②充填于气孔或空隙中的次生矿物，如火山岩气孔中充填的沸石、石英晶簇等。次生矿物还包括岩浆期后的气成矿物萤石、电气石等。
一些蚀变作用和交代作用常伴随矿化现象，因此，研究蚀变交代过程，对岩浆期后矿床的普查找矿意义重大。
(二)火成岩化学成分与矿物共生组合的关系
不同类型的火成岩其矿物成分不同，不同造岩矿物之间构成有规律的共生组合。其组合一方面与岩石形成当时的温度、压力等有关，另一方面主要是取决于岩石的化学成分。化学成分中尤以SiO2、K2O+Na2O、Al2O3的含量影响最大。
1.SiO2含量对矿物共生组合的影响
前已述及，SiO2是火成岩中含量最高的氧化物，与其他氧化物结合可形成各类硅酸盐矿物。当SiO2含量过剩(过饱和)时，其会从硅酸盐熔体中游离出来结晶成石英，故石英的出现是火成岩SiO2过饱和的标志。当SiO2含量不足(不饱和)时，岩石中出现SiO2不饱和的矿物，无石英生成，因为当这些矿物形成后，若岩浆中有多余的SiO2时，二者将发生反应生成其他矿物，例如:
岩石学
人们习惯于把火成岩中可以与石英共生的硅酸盐矿物称为SiO2饱和矿物(或硅酸饱和矿物)，如辉石、角闪石、斜长石、碱性长石、云母等;将不与石英共生的硅酸盐矿物称为SiO2不饱和矿物(或硅酸不饱和矿物)，如镁橄榄石、似长石(霞石、白榴石)、黄长石、黑榴石等;石英则称为硅酸过饱和矿物。
如前节所述，火成岩中各主要氧化物随SiO2含量变化而呈现规律性变化。反映在矿物成分上就是随着SiO2含量的增加，岩石中铁镁矿物由多到少，矿物种类从橄榄石、辉石变化到角闪石、黑云母;硅铝矿物则由无到有，或由少到多，矿物种类由富Ca向富Na、K、Si的方向演变(图2－16)。
图2－16 火成岩矿物组合变化图(Adams，1956)
2.碱质含量对矿物共生组合的影响
不同碱质(K2O+Na2O)含量的火成岩中矿物组合也有很大的差别。如前所述，根据里特曼指数σ的大小，火成岩可划分成钙碱性岩、碱性岩和过碱性岩类型。不同类型岩石的矿物组合明显不同，σ＜3.3的钙碱性岩石中不出现似长石、黑榴石和碱性暗色矿物(霓石、钠闪石、星叶石等)，出现长石、石英和普通辉石、透辉石、斜方辉石和普通角闪石等。σ＞9的过碱性岩中，常出现似长石和碱性暗色矿物(霓石、霓辉石、钠闪石、星叶石、富铁云母等)，长石主要为碱性长石，黑榴石也较常见，不见斜方辉石和石英。σ=3.3～9的碱性岩石中，常见的是碱性长石和碱性暗色矿物，可以出现石英、似长石(二者不共生)和除钠长石以外的斜长石。
3.Al2O3含量对矿物组合的影响
根据Al2O3与Na2O+K2O、CaO含量之间关系，火成岩可分为过碱质、过铝质和偏铝质三种类型，不同类型岩石有其特征的矿物组合。过碱质岩石中出现碱性长石、似长石和碱性暗色矿物;过铝质岩石中除长石、石英、黑云母外，还出现白云母、黄玉、电气石、锰铝－铁铝榴石、刚玉、红柱石、矽线石、堇青石等富铝的矿物;偏铝质岩石中则不出现上述岩石中的似长石、碱性暗色矿物以及大多数富铝矿物，而出现长石、石英、普通角闪石、普通辉石、透辉石和黑云母等。
(三)火成岩形成条件与矿物共生组合的关系
火成岩形成的物理化学环境对矿物组合也有重要影响。岩浆在地壳较深部位冷却时，处于温度缓慢下降、压力相对高的环境中，结晶时间充足。开始晶出的矿物，有些可能是高温型(如透长石、β－石英)，但随着温度缓慢下降，早形成的高温矿物不再稳定，逐渐转变为适应低温环境的稳定矿物。如透长石转变为正长石、β－石英转变为α－石英。因此，深成岩以出现低温矿物组合为代表。岩浆喷出地表时，环境由地下的高温高压急速变成常温常压，岩浆快速冷却来不及结晶而形成大量玻璃质，或生成颗粒细小的高温矿物组合岩石。同时，岩浆喷出地表带出的先前在地下结晶的高温矿物，也来不及转变成较低温矿物，仍保留着高温矿物的结构。因此，火山岩的矿物组合以高温矿物、细粒矿物和玻璃质为特征。此外，地下深部高温、高压环境，因大量挥发分参与结晶会形成含挥发分的原生矿物。喷出地表的岩浆，因挥发分大量散失，很难结晶出含水矿物，即便是岩浆在地下深处结晶出的诸如角闪石、黑云母等含水矿物，也因其被岩浆携带至地表发生氧化、脱水而分解或部分分解，转变成磁铁矿、赤铁矿等其他矿物，使原有矿物全部或在边部呈现黑色、褐色，这一情况称为暗化现象。
『柒』 燕山期对流地幔注入大陆，华北东部克拉通陆壳的改造和燕山期造山带陆壳的形成
(一)燕山期对流地幔注入华北东部大陆
华北地台东部侏罗纪-白垩纪突发性强烈火成岩活动的发育表明，燕山期对流地幔注入大陆，包括大量热能的注入和物质的注入。
1.巨大的热通量(heatfluxes)从对流地幔注入大陆
众所周知，华北地台在燕山期突发性强烈火成岩活动之前是典型的克拉通，古生代华北地台金伯利岩岩浆捎带的上地幔橄榄岩包体的金刚石中矿物包裹体的热力学计算获得的地温相当于地表热流值为40～45mW/m2(郑建平，1999)。与华北地台东部相对照，鄂尔多斯黄土高原在燕山期和喜马拉雅期一直保持克拉通构造上的稳定，可作为地台东部燕山期大陆“活化”开始前的一个参照。基于上海奉贤-阿拉善左旗地学断面(国家地震局地学断面编委会，1992)提供的在鄂尔多斯中心部位的定边一带的3个热流值，计算获得的平均值为44mW/m2，它与金伯利岩喷发时反演的热流值是符合的。这样，我们可用Wyllie(1997)提供的陆壳岩石(r-T1-υ-H2O)熔融相平衡图解(图2-64)进行讨论。燕山期火成岩主要是偏铝的Ⅰ型花岗岩类，需要达到使黑云母消失(Bi-out)(图2-64)的温压条件。英云闪长岩(T1)在1GPa压力(约35km深度)下脱水熔融实验(Rutter ＆ Wyllie，1998)表明，固相线温度为825℃，在850℃～900℃温度间隔内Bi-out过程中局部岩浆数量增加快，900℃时熔浆数量可达22%，留下难熔残余为:Opx+Hb+Ga+Pl+Qz+Mt+Sp。冀东太古宙片麻岩中分布广泛的黑云母片麻岩在1GPa压力下脱水熔融实验(吴宗絮等，1995)表明，固相线温度为812℃，Bi-out为837℃，Hb-out为887℃，在812～950℃间隔内形成水不饱和的花岗质岩浆，数量为20%～30%，留下的难熔残余为Opx+Cpx+Ga+Pl+Qz+Ru(金红石)。如果我们取大陆地壳的平均值为35～40km厚和地盾地温作为燕山期最初抬升的条件的话，那么，由图6-7可知，在1GPa条件(约35km厚的壳底部)下，最初始的温度大约400℃，要使陆壳底部岩石发生≥20%的局部熔浆(partial melt)的话，必需升高温度达850℃左右，亦即要使陆壳底部岩石加热升温约450℃才行(即初始400℃+升温450℃=850℃)。
图2-64 花岗岩-英云闪长岩-辉长岩-H2O熔融相平衡图解(据Wyllie，1997)
从对流地幔中分离出来的玄武质岩浆底侵注入壳底是对陆壳加热诱发局部熔融发生的最好机制(例如，Fyfe et al.，1973;Bergantz，1989)。Bergantz(1989)对玄武质岩浆底侵和陆壳岩石局部熔融进行了1GPa条件下一维定量热模拟，注入的玄武质岩浆温度为1250℃，玄武质岩浆通过自身的冷却和结晶对上覆的地壳加热，岩浆房内以及热能的传送到陆壳岩石内均考虑为传导(conctive)方式，模拟的结果为，当围岩陆壳岩石为英云闪长岩(T1)及其初始温度为700℃时，产生的总局部熔浆与玄武岩岩浆结晶总量的比值为0.4，例如，要产生500km3的熔浆，需要有1250km3的底侵玄武质岩浆的全部冷却固结，进而500km3的熔浆总量中只有25%才能有效地从熔融区中分离出来，上升形成火成岩，所以，只能形成125km3的火成岩，亦就是说，形成火成岩总量与底侵玄武岩岩浆的比值为0.1。由上，如果我们只考虑厚度(即一维模型)，则形成5km厚的花岗岩岩基，就需要50km厚的玄武质岩浆的底侵，需要含局部熔浆总量为20km的陆壳熔融区供给。与模型上地幔岩(pyrolite)类似的二辉橄榄岩(KLB-1，和HK66)(其化学组成见表2-11)的1GPa压力下熔融实验(Hirose ＆ Kushiro，1993)获得1250℃玄武质岩浆的熔融程度分别为6.5%和17.9%。这样，如要取熔融程度为10%的话，则生成50km厚的玄武质岩浆需要由500km厚的对流地幔的供给(如要全部局部熔浆均可分离出来的话)。需要注意的是，上述Bergantz(1989)定量热模拟是假设陆壳的初始温度已高达700℃的情况，从燕山期开始时，约35km深度处的陆壳岩石只有400℃，亦就是说，必须先把400℃陆壳升高300℃，才能达到700℃的初始条件，可以想象，要使400℃的陆壳加热升温达700℃，这需要由大量的底侵玄武质岩浆的注入大陆才能实现。另外，燕山期还有不少玄武质火山岩和辉长岩(υ)形成，亦就是说，除了要加热使陆壳熔融，底侵玄武质岩浆全部固结(这部分岩浆无法上升形成火成岩体)之外，对流地幔中还分离出不少可上升达浅部形成火成岩体的玄武质岩浆。虽然，我们还要结合本区燕山期火成岩的性质进行定量热模拟，但是，Bergantz(1989)的热模拟已可作为一种指导(guide)，告诉我们燕山期火成岩活动需要极大量的底侵玄武岩岩浆的供给，从一维角度来看，可能需要有深度直达670km的界面的对流地幔的供给，棋盘岩发现的MgO含量高达19%的玄武科马提岩质岩浆可能是由这样的过渡地幔中分离出来的，它对上述热模拟定量结果提供了一种支持。显然，原来冷的克拉通岩石圈是无法提供如此大量的热通量注入陆壳的，或者说，从热通量的角度来考察，底侵的玄武质岩浆决不会从冷的岩石圈地幔中分离出来的，而必定是从下伏的热的对流地幔(或软流圈)中分离出来的。
2.对流地幔分离出来的物质(玄武质岩浆)注入大陆
注入大陆的对流地幔物质是指从分离出来的玄武质岩浆，常称为初生(或新生)陆壳物质(juvenile crustal materials)。关于燕山期玄武质岩浆的起源常有两种不同看法，或认为来自软流圈(对流地幔)，或认为来自岩石圈地幔。
从岩石学组成来看，岩石圈地幔是对流地幔(或软流圈)分离出玄武质岩浆后的难熔残余，主要为方辉橄榄岩，软流圈则为丰满的地幔橄榄岩(含有丰富的玄武质组成)，接近原始地幔组成，Ringwood(1975)据此提出Pyrolite模型(上地幔岩模型)(表2-10)。
表2-10 Pyrolite模型
(据Ringwood，1975)
表2-10中，大洋拉斑玄武岩与方辉橄榄岩是一对互补产物，17%的拉斑玄武岩加83%的方辉橄榄岩就代表在互补物分离之前母体的组成，Ringwood用辉石和橄榄石的英文名称的字头(pyr和ol)加上构成岩石时常用的字尾(ite)构成一个人造的岩石，叫pyrolite，以代表原始上地幔岩石学组成的一个模型，它与自然界二辉橄榄岩组成大体相当。反过来说，当从pyrolite中分离出17%的玄武质岩浆之后，留下的83%为难熔的方辉橄榄岩残余。与pyrolite相比，残留的harz.富MgO，贫Al2O3，CaO，和FeO，表2-10中FeO重量百分数几乎无变化，但MgO与FeO比值变化大，表明FeO的相对贫乏，而MgO的相对富化。Al2O3，CaO，和FeO为易熔组分，进入玄武质岩浆分离走了，MgO为难熔组分，留在橄榄岩中，它与残余中Cpx消失和橄榄岩中FeO组分增高的矿物学特征相适应。可以推测，如要方辉橄榄岩再一次遭受熔融作用，那么，发生熔融作用的温度一定更高，熔出的熔浆组成一定不同于常见的玄武质岩浆(如表2-10)。
高温高压下，二辉橄榄岩和方辉橄榄岩局部熔融熔出的岩浆组成是不同的(Hirose＆ Kushiro，1993;Kushiro，1990;Falloon ＆ Danyushevsky，2000)，选择典型样品列于表2-11，某些氧化物关系见图2-62。
表2-11 橄榄岩源岩及选择的一些熔浆组成(均无水100%)
从表2-11与表2-10对比，以及图2-65可以看出，高温高压实验所用的二辉橄榄岩和方辉橄榄岩源岩与Ringwood模型中pyrolite和方辉橄榄岩是类似的，相对于二辉橄榄岩来说，方辉橄榄岩富MgO，贫Al2O3，FeO*，CaO，亦就是说，它是原始地幔橄榄岩分离出玄武质低熔组分后的难熔残余。由表2-10和图2-62可以看出:①在相同的高MgO含量时，方辉橄榄岩熔出的熔浆SiO2更高;在相同SiO2含量条件下，方辉橄榄岩熔出的熔浆MgO更高(图2-65a);②在相同的高MgO含量条件下，方辉橄榄岩熔出的熔浆贫FeO，相同FeO含量下，方辉橄榄岩熔出的熔浆富MgO(图2-65b);③总体上看，方辉橄榄岩熔出的熔浆Al2O3低，在相同的高MgO条件下，方辉橄榄岩熔出的熔浆则Al2O3高(图2-65c);④在相同的高MgO条件下，方辉橄榄岩熔出的熔浆低CaO(图2-65d)。
自然界大部分岩浆来自二辉橄榄岩源，少量(如汤加，Troodos)则来自方辉橄榄岩源(表2-12，图2-66)。
华北东部燕山期火山岩中MgO高的玄武岩和玄武安山岩大多亦是进化的岩浆，如要加入被分离走的橄榄石(Ol)的话，则MgO会更高，同时SiO2会有所减小，它们将会进入源区为二辉橄榄岩局部熔浆范围内(表2-12，图2-66);一部分辉长岩侵入体的化学组成，如棋盘岩和上庄岩体，MgO高相当于科马提岩-苦橄岩-玄武岩的组成，在图2-66上均相当于源区为二辉橄榄岩的局部熔融范围(表2-12，图2-66)。这样，燕山期幔源岩浆的源区不是太古宙形成的以方辉橄榄岩为主的地幔岩石圈，而是丰满的二辉橄榄岩为主的对流地幔或软流圈。
图2-65 方辉橄榄岩与二辉橄榄岩熔融产物对比图
表2-12 选择的自然界幔源岩浆组成(均为无水100%，P2O5未列)
续表
注:*郁建华等，1989，北京地区辉长岩深成作用及成因演化。
图2-66 方辉橄榄岩与二辉橄榄岩熔融产物对比图
这样，从燕山期火成岩形成的需要的热通量和幔源岩浆的注入的2个侧面来看，燕山期对流地幔注入大陆是导致华北克拉通东部大陆“活化”的驱动力。
(二)壳幔相互作用:混合、分异、与拆沉
1.岩浆混合作用:壳幔物质混合的主要机制
1)Zartman ＆ Doe(1981)铅(探测)构造(plumbotectonics)模型(第2版)。铅(探测)构造模型试图模拟在主要地球储库之间U、Th和Pb的地球化学习性(图2-67，图2-68)。图2-67强调了地球内的三个长期演化的储库，包括地幔，下地壳和上地壳，和一个短期的在造山带环境下上述三个储库输入造山带的动力学相互作用(dynamic interac-tion)形成造山带新地壳(图2-67的带黑点的柱)。图2-68展示铅(探测)构造模型预示的三个长期储库以及它们短期输入形成的造山带新储库的铅同位素演化曲线，它们的端点表示这4个储库的现代同位素组成。这样，造山带新地壳的形成是地幔注入的物质与先存的老的陆壳物质混合的结果。
图2-67 地幔、上地壳、下地壳和造山带铅同位素分配储库结构图(据Zartman ＆ Doe，1981)
图2-68 铅构造模型试图模拟在主要地球储库之间U、Th和Pb的地球化学习性图解(据Zartman ＆ Doe，1981)
2)燕山造山带火成岩的铅同位素组成。燕山造山带火成岩(SiO2从45%左右的玄武质→76%左右的花岗质)的铅同位素组成明显低于Zartman ＆ Doe铅(探测)构造模型(1981)中的4个储库的值(参见图2-67～图2-69)和表2-13，图2-69。
图2-69 铅构造模型试图模拟在主要地球储库之间U、Th和Pb的地球化学习性图解(据Zartman ＆ Doe，1981)
表2-13 Pb同位素组成的比较
由表2-13和图2-70看出，①幔源的棋盘岩科马提岩质辉长岩的Pb同位素显然低于地幔，甚至低于全球下地壳，表明它已与同位素低值的下地壳发生强烈的混合作用;②燕山期火成岩Pb同位素的低值区低于冀北-辽北亚省下地壳，表明下地壳源区的Pb同位素比张理刚(1995)提出的要更低，我们用变质结晶基底9个长石铅计算的Pb同位素值代表下地壳源区可能更合理些，它与地幔Pb同位素的混合可更合理地解释火成岩Pb同位素组成;③燕山期火成岩Pb同位素组成记录了岩浆混合作用和壳幔混合作用;④总的看来，似有一种趋势，J1和J2的火成岩Pb同位素组成大于J3和K1的火成岩，有可能预示J1和J2时期的岩浆活动中地幔的贡献大于J3和K1，这似乎与热模拟预示的J3和K1时地壳温度高于J1和J2符合，亦与J3和K1时期岩石圈面型拆沉和面型岩浆活动符合，由于资料仍较少(共23件)，还需要进一步积累更多资料，检验和修正;⑤对于同一个时期形成的杂岩体来说，例如薛家石梁杂岩体(除黑熊山花岗岩之外)，同位素演化范围更窄，大体上随SiO2含量的升高，同位素的比值略减小(图2-69)，符合了岩浆混合作用和壳幔物质混合作用的模型，黑熊山花岗岩的明显低值以及Sr同位素的明显的高值表明它来自另一个源区。
图2-70 燕山地区中生代火成岩SiO2-206Pb/204Pb图解(A)、SiO2-208Pb/204Pb图解(B)
图2-71 岩浆混合作用时元素和同位素交换的解耦(据Leisher，1990)
3)岩浆混合作用时元素与同位素交换的解耦。Leisher(1990)通过基性和酸性熔浆的实验研究，提出岩浆混合作用时元素和同位素交换的解耦，同位素的交换和混合明显地快于元素的交换和混合(图2-71)。从图2-71还可看出，不相容元素的交换和混合作用要快于主元素。
Leisher(1990)的实验成果是与自然界见到的岩浆混合作用符合的。在野外露头上我们见到的花岗质岩浆中含有微晶闪长岩的包体(MME)是岩浆混合作用的直接证据(莫宣学等，2002)。一般认为，含微晶闪长岩包体的花岗质侵入体的岩浆混合作用主要是物理混合(或机械混合)(magma mingling)，还没达到化学混合，它对包体与寄主岩之间的岩石学组成和主元素化学组成而言是正确的，因为包体常常是闪长质的，寄主岩则常常是花岗质的，同时，常常认为微晶(粒)包体预示它的快速冷凝固结，阻止化学混合作用的进行。然而，包体与寄主岩的Sm-Nd同位素研究(Pin，1991)表明，包体的εNd(－3.2→－4.1)和寄主花岗岩的εNd(－3.5→－4.1)是相同的，它表明同位素已经充分的交换并达到平衡，虽然主元素和岩石学组成仍有大的差异。由图2-67和图2-68可以看出，虽然SiO2含量的变化大，但Pb，Sr，Nd同位素变化范围小，特别是同一个杂岩体中岩浆混合作用形成的组合(如薛家石梁杂岩体中υ-δ-η-ξ)中同位素组成已有较充分的交换和混合，薛家石梁杂岩体中的黑熊山花岗岩不属于上述混合作用的系统，所以其同位素组成有大的差异(图2-72)。这样，火成岩中同位素不仅可以示踪和识别不同的岩浆系统，而且可以更好的示踪壳幔物质混合作用及其相对贡献的大小。如果基于质量平衡原理，按表6-7对206Pb/204Pb分别在地幔为18.01和17.616，陆壳结晶基底14.576，和燕山期火成岩16.369，则地幔与下地壳在混合作用时的分别贡献为52.2%和47.8%。
图2-72 薛家石梁环状岩体铅同位素组成图解
如果用143Nd/144Nd来估算，取宁芜地区燕山期蒋庙和阳湖堂辉长岩2个平均值0.5126385(εNd=+1.35)，华北陆壳结晶基底值0.511796(郭敬辉，2000)和燕山期火成岩0.511927，则估算的地幔和下地壳的分别贡献为15.5%和84.6%。
4)燕山期火成岩构造组合。燕山期火成岩组合为，玄武岩-钾质粗面玄武岩-钾玄岩-安粗岩-粗面岩和粗面英安岩-流纹岩的火山岩组合，和相应的辉长岩-二长辉长岩-二长闪长岩-二长岩-正长岩和石英二长岩-花岗岩的侵入岩组合(见本章第三节)。以火山岩组合为例，如图2-73所示，按SiO2含量分类，构成正态分布(或单峰式)，以安山质(A)为峰值，如按TAS分类，研究区的安山质(A)应为粗安岩;按SiO2-K2O分类，研究区火山岩主要为HKCA系列，这样，它们是典型的造山带火成岩组合;与安第斯比较则缺乏MKCA系列及其重要成员的安山岩(闪长岩)和英安岩(花岗闪长岩)而发育HKCA系列的安粗岩(二长岩)和粗面岩(正长岩)和粗面英安岩(石英二长岩)(邓晋福，刘厚祥等，1996)。
上述火成岩构造组合表明，壳幔物质的混合作用形成以安粗岩为峰值的HKCA系列，如要说太古宙陆壳以T1T2G1G2构成的话，则燕山期由于对流地幔物质(以玄武质岩浆的形式)的注入，使当时的陆壳“基性化”了，因此，往SiO2含量偏低的安粗岩方向移动，同时又相对高K2O(高碱)了，这是因为太古宙陆壳的以MKCA系列为主为低K2O(低碱)的T1T2G1G2经历局部熔融分离出大量高K2O(高碱)的花岗质-正长岩质岩浆有关。从TAS图和SiO2-K2O(图2-46)关系看，从对流地幔中分离出来的原生岩浆或近似原生岩浆来说，它们属MKCA系列，K2O和碱并不高，因此，使整个火成岩系列主体进入HKCA系列，和K2O、碱的升高主要和陆壳熔融有关。
2.壳幔物质分异的主要机制:局部熔融与AFC模型
1)局部熔融作用(partialmelting):不管是原先存在的太古宙T1T2G1G2和基性镁铁质变质岩，还是在燕山期先期底侵玄武质岩浆或是与陆壳混合之后形成的长英质岩浆固结成岩之后，遭受后期的加热均可发生局部熔融作用，此过程在形成岩浆的同时必留下互补的难熔残余物，因此导致壳幔物质的分异和演化。如要局部熔融程度和源岩类似，但压力不同，则产生的岩浆及其残留物的矿物组合将是不同的。
图2-73 青藏高原大陆碰撞火山岩、东燕山期火山岩与安第斯弧火山岩的对比
图2-74为玄武质组成的角闪岩熔融相图及残余矿物组合(Wyllie et al.，1997);图2-75为长英质岩石局部熔融形成粗面质(正长岩质)岩浆的与液相线平衡的矿物组合，后者即残余矿物组合(Deng et al.，1998);图2-76为奥长花岗岩(T2)熔融实验时，与液相线平衡的矿物组合，后者同样代表残余矿物组合(Laan ＆ Wyllie，1992)。
图2-74～图2-76的一个共同特点为，当陆壳岩石发生局部熔融时留下的残余矿物组合，在较低压力下(<1.4～1.7GPa)为含斜长石的麻粒岩相组合，在高压下(≥1.4～1.7GPa，约≥55km厚的陆壳)则为无斜长石的榴辉岩相组合，温度相对低时常有Hb，温度高时则只有石榴子石。
2)同化分离作用(assimilation-fractionational crystallization，AFC)模型:底侵的高温玄武质岩浆在对陆壳加温、同化陆壳岩石或者与陆壳局部熔融产生岩浆混合时，玄武质岩浆自身必须要发生强烈的结晶作用。燕山期火成岩的化学特征，包括AFM图解(图2-47)，SiO2-K2O图解(图2-46)，TAS图解(图2-45)和SiO2-(Na2O+K2O-CaO)图解(图2-48)均表明AFC模型是岩浆分异演化的一个主要机制之一，即从对流地幔分离出的原生玄武质岩浆是富MgO、低钾、亚碱、CA(按Peacock的碱-钙指数的分类)的特征，而陆壳源的长英质岩浆相对富FeO、高钾、富碱、AC(按Peacock的碱-钙指数的分类)的特征，通过AFC机制形成了总体上FeO/MgO比值变化很少、高钾(HKCA)、相对富碱和AC(按Peacock的碱-钙指数的分类)的火成岩系列，同样可按图2-74～图2-76推测，结晶分离出来的矿物组合在低压下为麻粒岩相组合，高压下为榴辉岩相组合。
图2-74 角闪岩脱水熔融相图(据Wyllie等，1997)
图2-75 H2O不饱和的安山岩液相线面(据P.J.Wyllie，1976，转引自邓晋福，1987)
图2-76 H2O饱和的Nuke片麻岩液相线面(据P.J.Wyllie，1992)
由上，壳幔物质的相互作用是通过岩浆作用一方面发生物质的混合，另一方面发生物质的分异和分离。
3.玄武质榴辉岩诱发岩石圈拆沉作用
美国大陆动力学研究的国家计划(肖庆辉等，1993)指出，在岛弧中形成的新地壳物质，总体成分相当于玄武岩，但大陆的速度结构则表明，现存的陆壳平均成分更接近于安山岩。也许在过去已在岛弧中形成了稍偏硅铝质的地壳，也许是某些作用长期以来把高密度的成分从陆壳中消除出去了。有关碰撞期间一部分大陆的下地壳和岩石圈地幔的拆沉作用，已被人们认为是值得研究的一种解释。如果这种作用经常出现，那么它对于了解大陆的结构和组成就具有深远意义。
上面的讨论告诉我们，①对流地幔物质的注入大陆是给大陆地壳增添大量物质的作用，由于大量玄武质岩浆的加入，必使原有的太古宙长英质陆壳向更镁铁质方向演化;②壳幔物质的混合同时，通过岩浆作用(如局部熔融作用，AFC过程等)壳幔物质又发生分异和分离，总体上分异出更富长英质的岩浆，分离留下更富镁铁质的矿物组合残余。但是，现今地球物理探测表明，鄂尔多斯块体陆壳平均vP为6.3km/s，太行和燕山造山带则为6.2～6.3km/s(嘉世旭，2003，未刊，内部报告)前者大体上相当于花岗闪长质(G1)组成(平均SiO2为65%)，后者大体上相当于花岗质(G2)组成(平均SiO2为70%)。鄂尔多斯块体自太古宙地壳形成以来一直保持构造上稳定的地台(克拉通)性质，因此，陆壳的组成可作为华北地台东部燕山期“大陆活化”之前原有陆壳组成的一个参照。燕山造山带基本上没有新生代玄武岩的分布，可总体上看作新生代没有地幔物质的添加作用，除了地壳厚度可能被减薄之外，其陆壳总体可看作为燕山运动造成的产物。因此可以看出，燕山运动使陆壳“酸性化”了，这与大量玄武质岩浆的添加使陆壳“基性化”形成了明显的矛盾，它必然要求相当部分的镁铁质又返回地幔，拆沉作用可能使最好的一种解释。
使陆壳“酸性化”的下地壳加地幔岩石圈的拆沉作用必须满足下述条件:①拆沉的下地壳(+地幔岩石圈)的密度必须大于下伏软流圈，对燕山造山带来说，还有一个特殊问题是，原有的太古宇岩石圈地幔是低密度的(以方辉橄榄岩为主);②拆沉的下地壳主体上应是镁铁质和超镁铁质的。能同时满足这两个条件的最佳对象是榴辉岩的大量出现。
前面已讨论，在壳幔相互作用的演化过程中，不断形成长英质岩浆的同时分离出镁铁质和超镁铁质矿物组合。造山带早期演化(J1和J2)的这部分镁铁质和超镁铁质残余矿物组分在J1晚期和J2晚期二次收缩构造驱动的陆壳加厚条件下，可能转变为榴辉岩相岩石，加上早期底侵的玄武质岩浆可能在使冷的陆壳加热过程中大部分冷却固结为岩石，他们在陆壳加厚条件下，亦会转化为榴辉岩相。J3和K11火成岩组合中出现没有或弱负Eu异常和高Sr/Y比值以及K2O高的安粗岩(二长岩)-粗面岩(正长岩，石英二长岩)表明，那时已是加厚的陆壳(≥55km)，底侵的玄武质岩浆的固结和分离出来的残余矿物组合应为榴辉岩相岩石。J1和J2的线性分布的火成岩和J3和K11面型分布火成岩表明，J3开始时堆积在加厚下地壳和岩石圈地幔内部的榴辉岩相岩石的体积已足够大(或已达一个临界值)，致使岩石圈发生大面积的拆沉作用。
(三)燕山期造山带陆壳的形成
现今剥露于地表的岩石，除少量新生代沉积岩之外，绝大部分是燕山期形成的火成岩(包括火山岩和侵入岩)、变质岩和沉积岩，以及卷入燕山运动的被改造的前侏罗纪岩石。不少大的以花岗质为主的岩基，如八达岭、大河南岩基，是剥露于地表的上地壳中下部的代表。云蒙山花岗质片麻岩岩基，沙坨子眼球状片麻岩体，五道河花岗质眼球状片麻岩体，石城闪长质片麻岩体，长园闪长质片麻岩体和源岩为长城-蓟县系的四合堂角闪岩相变质岩则是剥露于地表的造山带根带岩石，代表造山带中地壳岩石(Davis et al.，2001;北京市地质志，1991)。Davis et al.(2001)指出，奇怪的是闪长质片麻岩侵入体的锆石不含继承组分，但是，与新生的闪长质岩不同的是，花岗质片麻岩和眼球状片麻岩的锆石则含有继承组分，其上交点范围为1.7～2.4Ga。本项目的SHRIMP研究支持Davis等的结论，闪长质片麻岩侵入体无继承性锆石，而云蒙山花岗质片麻岩侵入体则有继承锆石(2.4Ga)。除此之外，新生代汉诺贝玄武岩中发现的麻粒岩相斜长辉石岩包体(120～140Ma)，和榴辉岩相石榴辉石岩(具堆晶结构)(樊祺诚等，1998，2001)，表明J-K时期确有新形成的下地壳麻粒岩相岩石和在山根带形成的榴辉岩相岩石，支持了燕山期曾有一个加厚陆壳和玄武质岩浆底侵的模型(邓晋福等，1996)。岩浆上升经过了太古宙和古元古代盖层。表明它们显然是从地幔中分出的新生物。
燕山期造山带陆壳的形成过程大致可归纳如下:
1)对流地幔的热和物质的注入，是大陆地壳生长和古老陆壳强烈改造的驱动力，注入的机制主要是幔源玄武质岩浆的底侵作用(basalticmagmaunderplating)。
2)大陆地壳的形成和改造主要通过火成作用实现，岩浆的形成和演化、运移和定位是大陆地壳生长的基本过程。
3)壳幔相互作用是大陆形成和演化的关键机制，岩浆过程是实施壳幔相互作用的主要机制，它包括壳幔岩浆的混合作用，岩浆的分异及其与残余的镁铁质矿物组合的分离过程，以及镁铁质和超镁铁质榴辉岩相岩石的拆沉过程。
4)造山带收缩构造驱动的加厚陆壳的形成是形成大量榴辉岩相岩石的必需，是下地壳和岩石圈地幔拆沉的必需。
5)大陆地壳的形成和改造，不仅有对流地幔物质的添加作用，同时亦有地壳和岩石圈物质的返回对流地幔，这两种机制是实现大陆地壳成熟化，最终形成以花岗质为主的成熟陆壳的必需。
6)岩浆构造事件序列以及火成岩构造组合是追踪对流地幔注入大陆直至形成成熟陆壳过程的关键记录。
7)Pb同位素和Nd-Sr同位素组成的联合约束是大陆地壳生长改造的最好示踪记录，它不仅可能鉴别完全新生的陆壳(以太古宙TTG为代表)，还是对流地幔再次注入的改造型陆壳(以燕山期G为代表)，还可能反演改造型陆壳中地幔与原有古老陆壳的相对贡献大小。所以强调同位素组成是因为，①火成岩的岩石学与主元素，由于相当部分的玄武质榴辉岩的拆沉返回地幔，陆壳的组成总体上只相当于对流地幔注入改造后留下来的以长英质为主的那部分组成;②在壳幔相互作用中，同位素的交换速率比主元素、痕量元素快得多，可能代表壳幔物质混合作用中的相对贡献;③相对于主元素和痕量元素来说，壳幔物质的端元成员的同位素比值的参考值容易确定，因而计算它们的相对贡献更为合理，可操作性强。
『捌』 元古宙祁连古陆块大规模岩浆事件及熔离成矿
（一）元古宙祁连古陆块大规模岩浆事件
金川矿床以其巨大的Cu-Ni-PGM储量而著称于世，作为一个出露面积仅1.34km2，而镍金属储量达546×104t，岩体矿化率高达60%的独立超镁铁岩体，世界上绝无仅有。在如此矿化高度集中的成矿事实面前，该矿床深部熔离或者称为侵入前大规模的硫化物液相与硅酸盐熔体发生不混溶的机制是肯定存在的，最终成矿岩浆房其实就是矿浆房，是极度富集硫化物液相的熔体就位冷凝结晶成岩成矿的。这一认识已为中国学者早以认识并不断得以丰富（汤中立等，1995），汤中立院士提出的“小岩体成大矿”概念是这一成矿机理的深刻揭示。如此大规模的金属硫化物聚集，必然有更大规模的岩浆源提供才成为可能，这是岩浆Cu-Ni-PGE硫化物矿床学研究进入20世纪90年代后期，普遍开始关注的问题（Keays，1997；Pirajno，2002）。大量的硫来源于地幔还是地壳，硫同位素研究结果，答案是多样的，有幔源的认识，如金川等（汤中立等，1995；李文渊，1996），有壳源或壳源加入的，如Noril'sk等（Naldrett et al.，1996）。并已深刻认识到硫的饱和度是影响硫化物液相不混溶或熔离的主要原因（Naldrett，1989；Rad'Ko，1991；Brugmann et al.，1993；Keays，1995），而降低硫在岩浆中的饱和度的因素，地壳物质的加入、岩浆混合等是主要作用（Ripley and Li，2002；Ripley，1981；Naldrett et al.，1993，1989，1999；Lambert et al.，1991，1999，2000））。但可肯定金属镍是岩浆来源的，大量的硫也主要来源于岩浆，而岩浆主要是幔源派生的。岩浆要提供超规模的镍聚集，必然有超规模的岩浆存在才成为可能。在陆壳环境下，超常规模岩浆存在最为可能的形式就是与地幔柱有关的大火成岩省（LIPs）（Lightfoot et al.，1997；Keays，1997）。因为只有来源于核幔边界（CMB）之下“D”层（地震学术语，为核幔之间的热和化学作用带，源于核的热能的传输带，认为地幔柱起源于该带，Lay et al.，1998）的地幔柱才有可能使局部地区突然发生超常热事件并产生大规模岩浆作用形成LIPs，并提供大量的成矿组分S、Cu、Ni和PGE。
大陆环境LIPs存在的最重要特征是大陆溢流玄武岩（CFB），目前已确认有大面积CFB分布的LIPs有俄罗斯的西伯利亚台地（Noril'sk矿床）、美国的陆中裂谷系统（Mid-continent Rift System）（Duluth，Me11en矿床）、南非的卡罗欧（Karoo）火成岩省、印度的德干（Deccan）高原和中国南方的峨眉玄武岩等，Pirajno（2002）在其“Ore Deposits andmantle Plumes”一书中总结与地幔柱有关的大火成岩省主要表现特征为（图4-45）：“D”层来源的地幔柱上升至岩石圈底部，由于减压（decompression）蘑菇状（mushroom-shaped）地幔柱头部发生深部熔融，并吸入岩石圈碎片，向上通过地壳裂隙时致使岩浆过滤形成地壳高位岩浆储库；其中部分储库抵达地表喷发形成大陆溢流玄武岩（CFB）或火山岩，其余则就地（岩浆房）固结形成层状火成杂岩。初始直接派生于地幔柱轴部的熔体高Mg贡献于溢流玄武岩的早期相，晚期地幔柱熔体的运移可能侵蚀岩石圈的热边界层并携带岩石圈碎片返回到地幔。因此，与洋壳岩石圈相比，地幔柱熔体不相容元素富集。总之，地幔柱-岩石圈作用熔融形成的LIPs主要表现为空间上三位一体的特点：其一是CFB；其二是地壳高位层状火成岩侵入体（layered igneous intrusions）；其三是基性岩墙群（mafic dyke swarms）。三者相互联系，并互为因果，尤其是CFB与大规模的基性和超基性层状侵入体关系密切，后者是岩浆房或供给上覆玄武岩的通道，因此在宏观和直接表现上CFB是LIPs最典型的展现。
图4-45 地幔柱与岩石圈地幔、地壳底部的镁铁-超镁铁岩浆库及侵位的大陆溢流玄武岩（CFB）和相应的岩床杂岩相互作用示意图
由于CFB暴露于地表所以研究最为详尽。全球14个CFB统计结果，从太古宙（2772Ma）到现代（15Ma）均有发育，发育周期1～＜10Ma，典型的倾向于Fe富集，主要由大陆拉斑玄武岩组成，太古宙CFB的底部有科马提岩产出，元古宙CFB中有西澳的Bangemall盆地拉斑玄武岩省形成于1.6Ga，与金川的年龄极为接近。在时空关系上CFB与地壳的隆起和扩张有紧密联系，所以既可以形成于陆中裂谷系统（美国的MCR，位于苏比利尔湖和堪纳萨斯，著名的Duluth层状杂岩即位于其中，成因同于大面积分布的基性熔岩），也可产在大陆边缘火山裂谷（北大西洋火成岩省）。基性岩墙群与大陆裂解有关，其放射状的基性岩墙群可能反映了地幔柱的中心位置，单个岩墙宽从几米至200m，长几百米至1000km，基性岩墙群是大陆重建的标志遗迹。层状火成岩侵入体，实际是一种概括的说法，大火成岩省中的大量不同类型侵入体均可列入其范畴，它们是岩浆Cu-Ni-PGE矿床的主要载体。
根据金川含Ni-Cu-PGE超镁铁岩侵入体巨大的金属量聚集，推测其为地幔柱作用结果并为地质历史上热点的认识，假设中元古代早期祁连山古陆曾发育地幔柱作用形成的LIPs，必然会存在层状火成岩侵入体（金川侵入体可视为其中成矿岩体之一）和相应的CFB、基性岩墙群分布。
（1）分布于北祁连西段镜铁山微地块朱龙关群中的巨厚火山岩系（详见第三章），可认为是CFB的局部出露或残留。夏林圻等（1999，2000）已从火山岩岩石学和地球化学角度详细论证了其属于CFB的特征，此不再赘述。认为CFB必然应有广大面积的分布，现今发现面积有限，可解释为由于成岩后所在祁连古陆块裂解而肢解和再次拼合隆起而不同程度地遭受剥蚀所致。龙首山西段发现的长城纪变质基性火山岩系也应是这一时期CFB的出露，由于缺少高质量的年龄数据和详尽的火山岩岩石学研究，目前还仅属于推测，但藏布太、青石窑超镁铁岩很可能是科马提质或苦橄质熔岩，可与朱龙关群中的大陆拉斑玄武岩和碱性玄武岩构成完整的CFB组合。
（2）基性岩墙群目前在祁连山区域范围内尚未见报道，但镜铁山微地块朱龙关群中CFB分布区内发现的大量辉长岩脉，可视作基性岩墙群的局部出露，北祁连西段1∶5万区域地质调查中亦发现有大量不规则的基性岩脉。另外，在龙首山大量平行分布于中元古代白云岩中的辉长、辉绿岩脉也是值得重新认识其地质意义的基性岩墙。由于过去未从较大尺度上考察祁连山前寒武纪基性岩脉的时空分布及其地质意义，基性岩墙群是否存在和存在的特征是亟待调查研究的问题。
（3）祁连山层状火成岩侵入体是以往研究中最为关注的对象，但祁连山（包括龙首山）对镁铁-超镁铁岩侵入体的成矿调查，迄今仅发现金川大规模成矿岩体外，龙首山分布的10余处镁铁-超镁铁岩侵入体并无矿床发现，却在南祁连古元古界基底化龙群中，发现有拉水峡等小型岩浆Cu-Ni-PGE矿床（详见第四章），最为值得注意的是尽管金属储量不大，但却是全岩矿化，岩体即矿体，意义独特。如果从祁连山古陆块中元古代早期LIPs视角认识金川矿床的存在，拉水峡含矿岩体与金川岩体是相联系的，是属同一地幔柱作用在岩石圈底部部分熔融上侵熔离贯入成矿的结果。预测还可能有隐伏的层状侵入体存在，有发现新成矿岩体的地质前提。
笔者提出祁连山中元古代早期存在大规模的地幔柱作用的大火成岩省（LIPs）的设想，旨在探讨金川这一世界级矿床的物质来源与所代表的地质涵义，并试图解释其下地幔源巨量金属物质在地壳高位的聚集。由于这一课题所蕴涵内容的复杂性，仅提出框架性的推想，但存在一些不容回避的问题。首先是年龄问题，金川岩体年龄目前出现的争议（中元古代/新元古代），镜铁山微地块中CFB大跨度的同位素测试年龄，拉水峡岩体的成岩成矿年龄等，均制约着对整个岩浆-构造事件清晰把握；其次是，由于岩浆作用范围为后期造山带及其边缘地块地质范围，造山带急剧的多次构造变动对先期火成岩事件已改造得面目全非，进行配套恢复难度极大。但无论如何，重塑祁连山前寒武纪构造-岩浆事件，对系统认识金川矿床的成矿演化过程，进一步开展区域性找矿部署有重要的意义。
（二）金川超大型世界级矿床的硫化物深部熔离成矿
前已述及，硫在硅酸盐熔融体中的溶解程度主要取决于FeO的含量，其次为CaO、MgO和Na2O的含量。据戈德列夫斯基（1981）对诺尔斯克含矿岩石硅酸盐熔融实验，高温下（1400～1500℃）以悬浮状态存在硫化物在硅酸盐熔融体中含量可达15%以上。Ma-cLean（1968）发现体FeO-Fe3O4-SiO2-FeS系中存在大量不混溶相（图2-14）。他指出在结晶作用过程中，含有少量硫的均匀硅酸盐熔体可能使硫饱和而形成硫化物，从而导致熔离硫化物液相存在。结晶轨迹取决于原始化学成分和结晶过程中氧逸度的减低、衡定或增加。在FeS-FeO-Fe3O4-SiO2体系中，两种液相存在的最低温度是1140℃。在这种简单的基性岩浆中硫的溶解度可能约4%，硫化物、氧化物液相中的二氧化硅约1%。自然界实际硫的溶解度可高达15%（戈德列夫斯基，1981）。这种不混熔硫化液相的结晶在硅酸盐完全结晶后才开始。金川岩体的成因矿物学研究表明，铬铁矿中Ni含量亏损（Barnes et al.，1999）、橄榄石岩浆包裹体中少见硫化物子矿物（杨轩柱等，1991），不混熔作用于硅酸盐矿物结晶之前。
自然界玄武岩熔体仅含0.03%重量的硫（MacLean，1968），比FeO-Fe3O4-SiO2-FeS体系中最大值要低100倍，而自然界岩浆FeO含量更低，硅酸盐中大量CaO、MgO和Al2O3与FeO结合，使FeO浓度大大减低，从而导致熔融体中硫溶解度降低。一旦硫化物熔体从硅酸盐熔融体中熔离，富硫化物熔离熔体很容易形成块状矿石。Ni、Cu在结晶早期阶段，由于N-O键比Fe-O键要坚固，而取代Fe进入硅酸盐结构（斜方辉石和橄榄石）。硫化物熔体熔离之前，硅酸盐结晶程度愈高，氧逸度愈高，Ni更倾向于进入硅酸盐矿物，很少倾向于与硫化物熔体结合。因此，金川矿床高含矿率的特点，硫化物熔离前不会有橄榄石的明显结晶，否则难以形成富矿。
Naldrett（1989）认为不同矿床物质成分的差异，是由于硫化物分离并分支平衡的结果。溶于岩浆中的硫化物，由于镁铁-超镁铁岩浆对富硅物质的同化作用，而呈不混熔液态而熔离。不同比例的硫化物分离是不同程度同化作用的结果。硫化物液相突然熔离下沉的原因可能是由于富硅岩浆混合的结果。可见岩浆混合作用对岩浆硫化物矿床成矿作用的重要。Lambert et al.（1988，1989）运用稀土元素、Sm-Nd同位素和Re-Os同位素地球化学研究，结合野外、岩石学特点确定Stillwater杂岩至少存在有两种端元岩浆，一种是U型岩浆，为Wyoming太古宙准大陆岩石圈地幔部分熔融的玄武岩；另一种是A型岩浆，为上地壳混染的玄武质岩浆与下地壳镁铁-超镁铁岩部分熔融合成的拉斑玄武岩。认为Stillwater杂岩堆积相对较宽的γOs、εNd值，为不同岩浆地球化学混合，在整个成岩成矿岩浆房历史过程中作用的结果。国外其他岩浆硫化物矿床，也具这种特征，如Noril'sk的苦橄岩和富PGE的Ni-Cu硫化物矿石，γOs=+4.1～+14.2（Lambert et al.，1989）、Bushveld杂岩UG-2铬铁矿和梅基梅斯层，γOs=+33～+68（Hart and Kinloch，1989）、Sudbury火成杂岩，γOs=+322～+854（Walker et al.，1991）。这些数据暗示在每一阶段系统中都存在某种程度上的地壳物质加入。金川岩体εNd＜0的特征（李文渊等，2004），γOs=+9.1～+122.6（刘民武，2004）也反映了地壳物质加入的特征，它是金川岩体结晶前硫饱和致使硫化物液相发生大规模熔离或不混溶作用的主要原因。图4-46可示意其成矿过程。
古地幔柱在抵达岩石圈底部由于减压发生部分熔融产生大量的岩浆，岩浆中有岩石圈碎片的不断加入，因此不断上升的岩浆（1），成分会发生变化；上升至岩石圈地幔与地壳界线处形成规模宏大的岩浆房（或称之为岩浆储库），这时候的岩浆硫还是不饱和的，随着岩浆房地壳物质的带入混染和新的岩浆注入发生岩浆混合（2），促使岩浆房中的硫饱和，大规模持续发生硫化物液相与硅酸盐岩浆熔融体之间的不混溶（硫化物液相深部熔离）（3），Ni、Co、Cu、PGE等亲硫金属元素纷纷进入硫化物液相中，致使岩浆房上部的基性硅酸盐岩浆熔融体亏损S、Ni、Co、Cu、PGE；当亏损金属元素的岩浆上侵刺穿（熔蚀）地壳在地表喷发溢流形成基性熔岩（5）的同时，大量的基性岩墙群作为岩浆管道亦将保存下来，更为重要的是深部岩浆房中下部的含硫化物岩浆或直接为硫化物液相由于地壳的挤压作用而向上侵入和贯入，形成地壳高位含硫化物岩浆房（或矿浆房）而就位（4），这样的含硫化物岩浆房应该是多个，但可能规模大小和含硫化物的量有所不同，且每个岩浆房岩浆的注入和矿浆的灌入可能不止一次。最终高位岩浆房发生岩浆结晶堆积和硫化物液相位于岩浆房底部，形成层状侵入体。后期的构造变动使含矿层状侵入体空间位置发生变化，陡立或侧伏并剥蚀出露地表或位于地壳浅部，成为可利用的矿床或潜在的矿床。
图4-46 金川岩浆Cu-Ni-PGE矿床硫化物熔离成矿演化示意图
金川矿床作为世界超大型岩浆Cu-Ni-PGE矿床，其成矿背景和矿床成因一直是人们感兴趣的问题，国外岩浆Cu-Ni-PGE矿床研究者也表现出了极大的兴趣。但总体来说，对外介绍不够，Sudbury矿床发现100年来，新认识层出不穷，仅著名的“EconomicGeology”杂志，已先后于1971年、1990年、2000年和2002年出版了四期专辑，由矿床学研究引发出来的学术问题愈来愈广泛。由于金川矿床在中国岩浆Cu-Ni-PGE硫化物矿床中有举足轻重的地位，其成矿地质背景和矿床成因的认识，牵涉到整个中国同类矿床的研究。关于岩浆混合成矿作用的观点，关于地壳物质的加入等等，均是大家十分关注的问题。金川岩体岩石化学具科马提岩化学成分特征，矿石地球化学为拉斑质玄武岩矿床的特点，推测其中一端元岩浆为遭受地壳混染的科马提岩质岩浆（以岩体东段岩浆房为主），（Pt+Pd）/（Os+Ir+Ru）比值低，PGE含量本身低，仅有Ni-Cu矿石形成；西段岩浆房为拉斑质玄武岩浆，以形成Ni-Cu硫化物矿石和Pt、Pd富集体为特征。这种认识仅仅还是一种推测，尚需更细致的研究。
另外，金川矿床复合热液作用对成矿的贡献亦是显著的，而且可能存在多个活动阶段，特别是成岩成矿后期的历次构造改造中，各种高温复合热液的活动对富铜块状矿石的形成和PGE的局部富集是相当重要的，有必要进行深入研究。
『玖』 矿物化学特征及变质作用温压条件
麻粒岩捕虏体的主要组成矿物为辉石、斜长石和黑云母，其电子探针数据列于表4-17。数据主要由北京大学造山带与地壳演化教育部重点实验室电子探针室分析，分析仪器为EPM-810Q，测试条件：加速电压15kV，束流3.8nA。少量数据由中国科学院地质与地球物理研究所电子探针室分析，分析仪器及测试条件同上一节。
表4-17 基性麻粒岩捕虏体主要造岩矿物的平均化学成分（wB/%）及变质温度
续表
注：FeO*为全铁；（a）为中国科学院地质研究所电子探针室测试，其余由北京大学地质学系电子探针室分析；TWB为以Wood-Banno（1973）二辉石温度计计算的结果，TW为以Wells（1977）二辉石温度计计算的结果。
1.辉石
辉石是麻粒岩捕虏体中最重要的造岩矿物之一，据Deer（1978）分类方案，斜方辉石为紫苏辉石，而单斜辉石属普通辉石和次透辉石（图4-24）。
图4-24 麻粒岩捕虏体中辉石的Wo-Di-Hd-Fs-En成分图
（据Deer，1978）
图中黑圈代表单斜辉石，空圈代表斜方辉石；A—透辉石，B—次透辉石，C—顽透辉石，D—普通辉石，E—斜方辉石；
1、2分别为与贫钙辉石共存的麻粒岩和火成岩单斜辉石的趋势线（张儒媛等，1981）
（1）斜方辉石
麻粒岩捕虏体中斜方辉石的镁铁比值XMg=MgO/（MgO+FeO）为0.41～0.50，多数近于0.5。在Opx的MgO和FeO*含量图（图4-25A）中落于华北地体麻粒岩斜方辉石（张儒媛等，1981；沈其韩等，1992；贺高品等，1991；阎月华等，1988；靳是琴等，1986；王时麒等，1994）和汉诺坝捕虏体麻粒岩斜方辉石（陈绍海等，1998）之间，相对于前者富Mg贫Fe，而较后者富Fe贫Mg。有关研究表明变质岩中斜方辉石的CaO含量不会超过1.5%，且CaO含量的增加反映其形成的温度增高。MgO-Al2O3-SiO2体系的实验研究表明，斜方辉石的Al2O3含量随温度升高而增加（Boyd，1973），在不与石榴子石共生的斜方辉石中，Al2O3含量随压力而增加（Obata，1976）。本区麻粒岩捕虏体中斜方辉石的CaO含量均小于1.5%，为0.49%～1.31%，平均0.98%；Al2O3含量为0.47%～1.81%，多数小于1.5%。和华北前寒武纪麻粒岩中斜方辉石相比，显示了CaO高而Al2O3和Na2O低的特点（表4-18），反映其形成温度较高而压力稍低。
图4-25 麻粒岩捕虏体中辉石及黑云母的MgO-FeO*变异图
图中数据出处见文中叙述
表4-18 麻粒岩捕虏体与华北克拉通前寒武纪地体麻粒岩中辉石矿物平均成分（wB/%）比较
注：1）数据引自张儒媛等（1981），沈其韩等（1992），贺高品等（1991）；2）引自阎月华等（1988）；3）引自靳是琴等（1986）；4）引自王时麒等（1994）。
紫苏辉石在变质岩中开始形成的温度较高，一般认为它标志着麻粒岩相变质的开始。然而，由于麻粒岩捕虏体的原岩可能是辉长岩类岩石，因此有必要判别这些斜方辉石是早期岩浆成因的还是经后期麻粒岩相变质作用形成的。Dobretsov（1970）曾提出以下判别式来判断斜方辉石的成因（式中元素符号代表辉石结构式中的相应阳离子数，Fe3+校正采用郑巧荣（1983）的方法）：
D（x）=-13.5+59.6AlⅣ+16.6Fe3++21.2Fe2++15.9Mn-5.12Mg+0.9Na
若D（x）大于0，为麻粒岩相（无石榴子石）的辉石；若D（x）小于0，则为岩浆成因辉长岩类的辉石。本区捕虏体麻粒岩斜方辉石D（x）均大于0，表明为麻粒岩相变质成因。另外，Bhattacharyya（1971）曾据Al2O3和（MgO+FeO+Fe2O3）的含量来区分火成和变质成因的斜方辉石，认为（MgO+FeO+Fe2O3）+0.775Al2O3大于44.304者为变质成因，反之为岩浆成因，本区斜方辉石该值均大于44.304，进一步说明为麻粒岩相变质作用产物。
据Dobretsov等（1973）数理统计，斜方辉石的形成条件综合反映在其成分上，并提出了以下判别式：
D（x）=-4282+683Si+2192AlⅥ+2182Fe3++1455Mn+1442Mg+1427Ca+1700（Na+K）
认为D（x）大于0，属高温麻粒岩相；D（x）小于0，属角闪麻粒岩相。本区麻粒岩捕虏体斜方辉石D（x）值均大于0，显示其形成温度较高，已达高温麻粒岩相。
（2）单斜辉石
麻粒岩捕虏体中单斜辉石（表4-18）显示了富Mg、贫Fe、低Al、Ti、Na的特点。单斜辉石镁铁比值XMg为0.5～0.64。在Cpx的MgO-FeO*图（图4-25B）中，其MgO、FeO*含量明显不同于华北前寒武纪地体麻粒岩中单斜辉石，而与河北汉诺坝玄武岩中麻粒岩捕虏体单斜辉石相近（樊祺诚等，1996）。实验资料表明，单斜辉石的AlⅥ、Na含量一般随压力增大而增加（张儒媛等，1981；Wood，1974）；CaO含量可反映温度变化趋势，随温度降低而增高，而压力对它的影响在1000℃以下可忽略不计（MacGregor等，1976）；Ti的效应与Na正相反，随压力增大，钛辉石分子含量减少，当压力大于10kbar时，它们不出现（Wood，1974）。本区麻粒岩捕虏体中单斜辉石Al含量较低，其中AlⅥ为0.028～0.103，平均0.074，大于AlⅣ的平均含量0.019，且二者未见明显相关关系。CaO含量为17%～22.4%。Na的含量变化较大，从0.003～0.04，平均为0.021，Na和AlⅥ呈负相关关系。Ti的含量很低，平均为0.007。和华北地体麻粒岩中单斜辉石相比，本区单斜辉石CaO含量略高而AlⅥ、Na含量低（表4-19），表明它们形成的温度略高而压力低，但Ti的含量相近，可能又表示压力不会太低。按矿物化学成分计算端元分子，单斜辉石的硬玉分子（NaAl［Si2O6］）很低，最高也不超过2%，意味着其形成压力不会太高。总之，本区麻粒岩捕虏体中单斜辉石的化学成分特征表明其形成温度较高而压力相对较低，与斜方辉石成分反映结果相一致。
单斜辉石成分与寄主岩成因密切相关。Vejnar（1975）曾采用Ti-Al2O3和（Si+Al）-Al2O3变异图将单斜辉石分为变质成因和岩浆成因两大类。在这两种图解上，本区单斜辉石均落入变质成因区（图4-26）。以Dobretsov（1970）提出的判别单斜辉石成因类型的方法进行计算，本区麻粒岩捕虏体中单斜辉石亦为麻粒岩相变质成因。
图4-26 麻粒岩捕虏体单斜辉石的Ti-Al2O3和（Si+Al）-Al2O3图解
A—变质成；B—岩浆成因，黑点代表麻粒岩中单斜辉石，阴影部分代表堆晶岩中单斜辉石成分区
从图4-24中可以看出，单斜辉石基本上均沿Binns所确定的麻粒岩相单斜辉石趋势线附近分布（张儒媛等，1981）。图中共生二辉石投影点连线与Wo-En交于Wo70En30与Wo80En20之间，说明二辉石是平衡共生的（阎月华，1997），且与火成岩中二辉石的平衡共生相似（Deer，1978）。
2.黑云母
麻粒岩捕虏体中普遍含有黑云母，其含量1%～10%不等，大多呈棕红色（Ng），这是TiO2含量高的标志。Ti对云母，尤其对黑云母的颜色有强烈影响。一般TiO2含量越高，云母颜色越红（阎月华等，1988）。
黑云母的化学成分见表4-17。从表中可以看出，其成分上较富Mg，大多数Mg/Fe比值大于2，已经属于金云母范畴，而华北前寒武纪地体麻粒岩中不存在金云母，两类麻粒岩的云母成分有明显区别（图4-25C）。这些黑云母较山东金伯利岩中典型金云母（周作侠，1988）以及本区同一寄主岩中所含堆晶岩捕虏体中金云母贫镁富铁，而和澳大利亚昆士兰基性麻粒岩捕虏体中黑云母（Rudnick et al.，1987）相当。另外，本区麻粒岩捕虏体中黑云母还显示了Ti高、Mn低、AlⅥ和Si都较低的特点。TiO2含量为2.4%～5.2%，MnO为0～0.1%（多数小于0.05%），Al2O3为13.8%～15.2%，AlⅥ在以22个O为基础计算的化学式中为0～0.19，AlⅣ相对较高，AlⅣ/Si=0.419～0.469，AlⅣ与AlⅥ间未见明显相关关系。
黑云母的上述Ti高、Mn低、AlⅥ低的特点清楚地说明捕虏体变质作用已达麻粒岩相。Ti高、Mn低、AlⅥ低正是麻粒岩相黑云母的特点（阎月华等，1988）。许多研究者都证明了黑云母中Ti的含量是变质程度的函数，Ti含量随变质程度增高而增加（张儒媛等，1981）。角闪岩相黑云母中Ti一般小于0.3（以22个O计算），而麻粒岩相黑云母中Ti则较高（Guitdotti，1984）。本区麻粒岩捕虏体中黑云母Ti一般都大于0.3，最高者达0.564。黑云母的AlⅥ也是变质程度的函数，低级变质岩中黑云母AlⅥ高，变质程度高则AlⅥ低，麻粒岩相黑云母的AlⅥ通常小于0.55，非泥质麻粒岩相岩石的黑云母AlⅥ更低（Guitdotti，1984）。本区黑云母AlⅥ最高也只有0.19，远小于0.55。另外，在变质黑云母的TiO2-100Fe/（Fe+Mg）图解中（张儒媛等，1983），黑云母样品的成分点均落在了麻粒岩相区（图4-27）。因此，可以说本区捕虏体麻粒岩中黑云母具有麻粒岩相的特点，为麻粒岩相变质作用的产物。
图4-27 麻粒岩捕虏体中黑云母TiO2-100Fe/（Fe+Mg）图解
3.斜长石
斜长石是麻粒岩捕虏体中主要浅色造岩矿物。一般斜长石对于寄主岩的化学组分、共生矿物和变质作用等级都较敏感，区域变质作用过程中形成的斜长石，其An含量随温度升高而增加，角闪岩相斜长石的成分主要为拉长石或中长石范围，麻粒岩相者则可达培长石、拉长石范围（叶慧文，1986）。据斜长石化学成分计算其端元化学组成，本区麻粒岩捕虏体中斜长石An=42～74，分别为中长石、拉长石和培长石。
『一』 What conditions are ferroferric oxide, ferrous oxide, and iron oxide formed by reacting with air?
Feferric oxide is formed by the violent combustion of iron in the air. Iron is the result of self-oxidation of iron in the air. As for ferrous oxide, it should be produced by self-oxidation under conditions of insufficient air (I am not sure about this).
『二』 What does iron and oxygen react at high temperatures to form ferric oxide?
Iron and oxygen react at high temperatures to form ferric oxide. At high temperatures, iron burns in pure oxygen, reacts violently, and sparks are emitted, forming iron tetroxide. The chemical equation is as follows:
(2) What does currency circle feo mean? Extended readingUses of iron:
1. Used in pharmaceuticals, pesticides, powder metallurgy, and thermal hydrogen Generators, gel propellants, combustion activators, catalysts, water cleaning adsorbents, sintering activators, powder metallurgy products, various mechanical parts products, cemented carbide material products, etc.
2. Pure iron is used to make the cores of generators and motors, reduced iron powder is used in powder metallurgy, and steel is used to make machines and tools. In addition, iron and its compounds are also used to make magnets, medicines, inks, pigments, abrasives, etc.
3. Used as reducing agent. Used in iron salt preparation. Also used in the preparation of electronic components.
4. Used as nutritional supplement (iron fortifier).
『三』 How to distinguish between FeO, Fe2O3 and Fe3O4! What's the connection between them? Speed! Thanks!
Iron oxide (FeO) is a black powder that is unstable. When heated in the air, it can be oxidized into ferric oxide and react with acid (weakly oxidizing acid).
(3) What does feo in the currency circle mean? Extended reading:
Main purposes:
1. FeO:
p>It can be used as a pigment and is used in cosmetics and tattoo inks. Ferrous oxide is also used in porcelain making to give the glaze a green color. However, this substance is unstable and can easily be oxidized into ferric oxide.
2. Ferrous oxide:
It is used for coloring paint, rubber, plastics, construction, etc. It is an inorganic pigment and is used as an anti-rust pigment in the coating industry. Used as a colorant for rubber, artificial marble, and floor terrazzo; as a colorant and filler for plastics, asbestos, artificial leather, leather polishing paste, etc.; as a polishing agent for precision instruments and optical glass; and as a raw material for manufacturing magnetic material ferrite components, etc. .
Used in the electronic industry, communications equipment, televisions, computers and other magnetic materials as well as line output transformers, switching power supplies and ferrite cores with high U and high UQ
Used as analytical reagents, catalysts and polishing agents, and also used as ingredients for pigments;
Used for coloring the outer coatings of various tablets and pills
Used as magnetic materials and pigments And prepare reducing agents, polishing agents, catalysts, etc.; used for coloring tablet sugar coatings and capsules, etc.
Used as pigments for anti-rust paints. Because the mica iron oxide anti-rust paint made from this product has good water penetration resistance and excellent anti-rust performance, it can replace red lead
edible red pigment. In Japan, it is used in red bean rice and konjac flour foods. Used to identify bananas whose stem cuts have been treated with preservatives. In the United States, it is mostly used in cat food, dog food and packaging materials
Inorganic red pigments are mainly used for transparent coloring of coins, but also for the coloring of paints, inks and plastics
Widely used in paints , rubber, plastic cosmetics, construction fine grinding materials, precision hardware instruments, optical glass, enamel, cultural and educational supplies, leather, magnetic alloys and coloring of high-grade alloy steel; mainly used as magnetic materials, pigments, polishes, catalysts, etc., and also Used in the telecommunications and instrument industries; mainly used as magnetic materials, pigments, polishes, catalysts, etc., and also used as inorganic red pigments in the telecommunications and instrument industries.
3. Ferric oxide:
Ferric oxide is a commonly used magnetic material.
Specially pure ferric oxide is used as a raw material for audio tapes and telecommunications equipment.
Natural magnetite is a raw material for iron-making.
Used to make primer and topcoat.
Feferric oxide is the main raw material for the production of iron catalyst (a catalyst).
It is very hard and can be used as an abrasive. It has been widely used in the field of automobile braking, such as brake pads, brake shoes, etc.
Feferrous oxide has been recognized in the field of domestic welding materials. The production of welding rods and welding wires is still in its infancy, and the market prospects are very broad.
Feferrous oxide has shown good performance in sewage treatment due to its high specific gravity and strong magnetism.
Fe3O3 can also be used as pigments and polishes.
We can also use certain chemical reactions, such as using sodium nitrite, etc., to generate a dense layer of ferric oxide on the surface of steel to prevent or slow down the corrosion of steel, such as firearms, The surface of the saw blade and other surfaces turns blue or black. Commonly known as "roasted blue".
『四』Hegenshan type ophiolite type chromium ore
Hegenshan type chromium ore: produced in the Erenhot-Hegenshan-Heihe ophiolite belt, that is, the Siberian plate and Xilinhot -The collision zone of the Songnen microcontinent (belonging to the broad Sino-Korean plate), the ophiolite was formed in the Late Devonian, and its emplacement was approximately from the Late Carboniferous to the Early Permian. The rock mass is composed of high-aluminum mantle peridotite and heapite. Localized melting-modified pod-like chromite occurs in the dunite-clinogite complex phase in the central part. The industrial type of ore is mainly refractory grade. The deposit The size reaches medium. Crystallized differentiated chromium ore in pile crystal lithofacies has little industrial value.
1. Hegenshan Chromium Deposit, Xilinhot City, Inner Mongolia
(1) Overview
Hegenshan Chromium Deposit is located 9km northwest of Chaoke Wulasumu Township, Xilinhot City, Inner Mongolia , geographical coordinates: 116°18'00" east longitude, 45°50'00" north latitude. The deposit includes three mining areas: 3756, 620 and 733, of which the Hegoola 3756 mine isThe area is the largest, consisting of 186 ore body groups. By the end of 1993, the cumulative proven C+D grade reserves were 1.299 million tons. It is a medium-sized chromium ore deposit and the largest chromium ore deposit in North China. The average Cr2O3 content of the ore is 23.63% ~ 27.26%, and the ferrochrome ratio Cr2O3/
In 1954, the 126th team of the Northeast Geological Bureau discovered ultrabasic rock mass and many mineral spots such as chromium and nickel in the area. In 1956, the team discovered ore for the first time in the 3756 exploration well. In 1963 Submitted the "Final Exploration Report of Hegeola 3756 Chromium Deposit in Xilingol League, Inner Mongolia" and submitted chromium ore reserves of 1.255 million tons. Since then, from 1964 to 1985, many units have conducted general survey, exploration and scientific research on the deposit to varying degrees, which has greatly deepened the understanding of the mineralization of chromium ore in the area, but there has been no new breakthrough in prospecting.
In 1958, the Xilingrad League of Inner Mongolia established Factory 582 and began production. By 1962, a total of 20,000 tons of ore had been mined, and the surface ore bodies had been exhausted. In 1971, it was invested by the state and planned to mine deep ore bodies. However, due to some technical and economic problems, construction was stopped in 1986. In 1988, it was put back into production, mainly mining rich ores. In 1995, another well was opened to collect deep rich ores, with an annual production capacity of 30,000 tons. A ferrochromium alloy plant was built in the urban area. At present, a single furnace in the first phase is put into operation, with a daily output of 60# chromium. Ferrous alloy 25 t. Sold to Henan, Shanxi, Liaoning and other places, production of the final concentrate powder was discontinued due to price issues and losses.
(2) Regional geological overview
The Erenhot-Hegenshan-Heihe ophiolite belt, the geotectonic location is between the Siberian Craton and the North China Craton in the Tianshan-Xingmeng Orogenic system (Ren Jishun et al., 1999), that is, the East Ujimqin Banner-Boktu Fe-Mo-Sn-W-Cu-Pb-Zn-divided by Xu Zhigang et al. (2008) in the "China Mineralizing Zone Division Plan" Ag-Au-Cr III-level mineralization belt (III-48) (Vm-I). This ophiolite belt is a symbol of the collision between the Siberian Plate and the Xilinhot-Songnen Microcontinent (belonging to the Sino-Korean Plate in a broad sense). The collision began in the Late Devonian and continued into the Early Carboniferous, resulting in the formation of There are obvious differences in the biota on both sides before the Weixian Period of the Early Carboniferous (Guo Shengzhe, 1986), and the earliest mixed species appeared in the Late Devonian (Tang Kedong and Zhang Yunping, 1991). As a result, the Hegenshan ophiolite belt and ophiolite-type chromium ore were formed (Figure 3-37).
Figure 3-37 Schematic diagram of ophiolite distribution in central and eastern Inner Mongolia
1—Quaternary; 2—Jurassic; 3—Carboniferous-II Superimposed series; 4—Devonian; 5—Cambrian-Silurian; 6—Mesoproterozoic; 7—Archean; 8—ophiolite belt; 9—granite
(quoted from Baoyin Wuliji et al., 2009, slightly modified)
Hegenshan ophioliteThe ophiolite belt is a relatively well-preserved ophiolite belt in northeastern Inner Mongolia. It not only provides valuable information for us to trace and infer the history of the extinct ancient Mongolian ocean basin, but is also the most important chromium ore producing area in Inner Mongolia and even North China. The exposed ophiolite belt is about 120km long, 20km wide and covers an area of about 2158km2. The ophiolite rock combination is relatively complete: lherzolite, which is closest to the original mantle, clinopyroxene-dunite representing the depleted upper mantle; mafic accumulation complex at the bottom of the magma chamber; layers of the main body of the magma chamber Shape-massive gabbro; as well as the final derivatives of the magma chamber, plagioclase granite and diabase sheet dykes, metamorphic basic volcanic rocks, siliceous rocks and deep-sea sediment cover, which can be completely compared with typical snakes at home and abroad. Greenstone profile comparison. It has been identified that there are more than 20 ultrabasic rock bodies in the ophiolite belt, including Ayula Sea, Oyotingela, Chaoge Udel, Hegeaola, Hebai, and Bai from west to east. Rock bodies such as Yin Aola, Xiaobaliang, Chonggen Aola, and Wusnihei are distributed in an S shape. Except for Chaoge Uder and Chonggen Aola, where large areas of ultrabasic rock are exposed, the rest are blind rock bodies hidden under the Mesozoic and Cenozoic sediments. According to geophysical data, it is inferred that the Ayola Sea rock mass has the largest burial depth, reaching about 900m, and other rock masses have different burial depths. Among them, the Hegeola rock mass is the most important.
The oldest formation in the area is the Middle Devonian, exposed at the southwest end of the rock mass. It is a set of metabasic volcanic rocks, volcaniclastic rocks, tuff, green schist intercalated with thin layers of chert and crystal. limestone. The eastern end of the rock mass is in contact with the Upper Carboniferous strata, which are mainly acidic volcaniclastic rocks and tuffs. The north, west and south sides of the rock mass are unconformably covered by Lower Cretaceous mud conglomerate, sandy conglomerate and siltstone. The radiolarian siliceous rock collected from the heapite was identified by Wang Naiwen as the Late Devonian (Bao Peisheng et al., 1999). Its formation time should be earlier than the Devonian, and its emplacement time was approximately from the Late Carboniferous to the Early Second Age. Stacked Epochs (Inner Mongolia Geological Bureau, 1984).
Fold structures are only found in the Middle Devonian and Upper Carboniferous strata. The Middle Devonian strata form a monoclinic structure, and the Upper Carboniferous strata present a syncline with an axial direction nearly north-south (outside the area) and upwards to the north. The main fault structures are mainly normal faults, followed by faults of unknown nature. The faults trend northeast and nearly east-west, dipping to the northwest and south respectively, with medium dip angles. The emplacement of the Hegenshan ophiolite is controlled by two groups of nearly east-west and north-northeast fault zones before diagenesis.
(3) Characteristics of the Hegaola rock mass
The basic-ultrabasic rock mass of Hegaola is mainly composed of pile crystal complex and local melted (depleted) mantle rock composition. This rock mass is distributed in the middle section of the Ayura Sea Hegeola-Usnihe S-shaped rock mass group. The surface is oval, about 9km long from east to west, 6km wide from north to south, and covers an area of about 50km2 (Figure 3-38). The rock mass strikes 30° to the northeast, dips 120°, and has an inclination angle of 67° to 78°. According to drilling data from the North China Oilfield, ultrabasic rocks have been seen at a depth of more than 3,000 meters. Geophysical data estimate that the deepest thickness can reach 5800 meters (exchange with Inner Mongolia Geological Survey Institute, 2011). The rock mass is surrounded by extensive middle and newIt is covered by biogeographic strata and outcrops are extremely rare. It is connected to the Hebai rock mass to the east.
The cumulate rocks exposed on the surface are composed of strongly altered plagioclase, olivine (forsterite) and clinopyroxene (diopside); while the localized mantle rocks are composed of It is composed of about 90% deformed metamorphic clinolite and 10% dunite and basic rock dikes. Dunite lenses are distributed in groups and are in rapid change or transitional contact with the clinolite. The main rock-forming minerals are olivine (forsterite) and orthopyroxene (enstatite). The composition of the accessory mineral chromium spinel in various rocks does not change much (Bao Peisheng et al., 1999).
Four lithofacies zones are divided according to rock combinations:
1) Pile crystal complex lithofacies zone (Zone I): distributed on the northeastern edge of the rock mass and extending for thousands of meters , with a maximum width of more than 570 meters and an area of about 0.6km2, accounting for 2% of the total area of the rock mass. The bottom is composed of feldspar dunite, dunite and light-colored peridotite. The three often transition to each other. Chromium mineralization is common in feldspar dunite. The upper part is olivine gabbro, with a total thickness of about 40 meters.
2) Peridotite-dunite-layered basic rock lithofacies zone (II zone): distributed on the east side of the rock mass, extending in the northeast direction, with a large width change of 1000 ~2200m,
three rock types (clinolite), φ1 (dunite), and v (gabbro) frequently alternate, which is equivalent to the crust-mantle transition zone. This belt produces some small chromium deposits, such as Jidong Mine, 620 Mine, 820 Mine, etc. and several mineralization points.
3) Peridotite lithofacies zone (Zone III): Distributed on the west side of Zone II, it is mainly composed of clinopyroxenite and sporadic dunite. There is a small amount of chromium mineralization, such as B265 mineralization point.
4) Dunite-clinogite complex lithofacies zone (IV zone, Hegeola 3756): It is the most important lithofacies zone of the rock mass, distributed in the near middle of the rock mass. Towards the northeast, it is mainly composed of clinopyroxenite and elongated (plastically deformed) dunite lenses. The two alternately appear as striped complexes. The pyroxene content in clinopyroxenite is generally <17 %, forming a low pyroxene zone; the pyroxene content in the clinopyroxenite on both sides is generally about 20% to 25%. The dunite detached bodies are relatively dense and distributed in the northeast direction. The number of dikes in the zone is small and the alteration is strong. These characteristics show that the mantle rocks in this area have a high degree of partial melting and strong plastic deformation, and are typical locally melted mantle residue phase rocks. This belt contains 3756 medium-sized chromium deposits that are mainly of industrial value (Figure 3-38).
The rocks are generally strongly serpentinized, accompanied by silicification and carbonation, and mineral residual crystals are rare.
According to data from the Inner Mongolia Geological Survey Institute, the average rock chemical composition of clinopyroxenite is SiO2: 35.5% ~ 37.8%, TiO2: 0 ~ 0.3%, A12O3: 0.27% ~ 0.5%, Cr2O3: 0.05 %～0.17%, Fe2O3: 0.71%～3.18%, FeO: 1.63%～4.29%, MgO: 56.85%～59.28%, CaO: 0~0.13%, K2O: 0~0.01%, Na2O: 0~0.2%. This shows that clinopyroxenite has high magnesium, low titanium, and extremely low potassium and sodium. Among them, the m/f ratio of section 3756 is 10.56, which is the highest degree of basicity. The m/f ratio of Lot 41 is 9.4, indicating a low degree of basicity. The m/f ratio of the Jidong section is 10.08, which is between the above two. The MgO/SiO2 ratio of clinopyroxenite is inversely proportional to the pyroxene content and directly proportional to the olivine content.
Figure 3-38 Geological map and cross-section of the Hegenshan rock mass
(Adapted from the 109 Geological Team of Inner Mongolia Bureau of Geology and Mineral Resources, 1980)
p>1—Quaternary; 2—Lower Cretaceous mud conglomerate-siltstone; 3—Upper Carboniferous volcaniclastic rock—tuff; 4—Lower—Middle Devonian metabasic volcanic rock—tuff intercalation Chert crystalline limestone; 5—rhyolite porphyry; 6—granite porphyry; 7—duntite; 8—clinogite (low-glow harzburgite); 9—long dunite; 10— Peridotite; 11—gabbro; 12—rock Moho; 13—lithofacies zone boundaries and numbers; 14—measured-inferred geological boundary; 15—measured unconformity geological boundary; 16—contact boundary or geological occurrence ; 17—Flow surface and occurrence; 18—Large and small chromium deposits; 19—Chromium deposits; 20—Chromium deposit range; 21—Syncline axis
Dune petrochemical characteristics: Section 3756 The m/f and MgO/SiO2 values are 9.73 and 2.76; these two ratios are the lowest in section 41. Compared with ordinary dunite, ore-bearing dunite has a higher Mg content and a higher magnesium-silicon ratio, indicating that the formation of chromium ore is most closely related to magnesium-rich dunite.
Pile crystal rocks are characterized by light rare earth depletion and strong positive Eu anomalies; while local melting mantle rocks have V-shaped (or U-shaped) (3756 ore bodies) with relatively low total REE content. The clinopyroxenite) or pipe-shaped (dunite of the Jidong ore body) distribution pattern and the characteristics of negative Eu anomaly. It is basically the same as the rare earth characteristics of rocks in my country’s ophiolite chromium deposits (Figure 3-39, REE distribution patterns of various rocks in Hegenshan).
Figure 3-39 REE distribution patterns of various rocks in the rock mass
(According to Bao Peisheng et al., 1999)
(a ) REE types of mantle rocks: 1, 2, and 3 are clinolite in 3756 area; 4 and 5 are dunite in 3756 mining area; 6 is clinolite in Jidong mining area; 7 is dunite in Jidong mining area. ;
(b) REE type of heapite: 8 is feldspar-bearing dunite; 9 is peridotite; 10 is laminated gabbro; 11 is peridotite in Jidong mining area; 12 is the gabbro of the 3756 mining area
(4) Deposit characteristics
The Hegoola chromium deposit has two types of chromium ore: one is the rock produced in the ophiolite profile Below the Moho, there are localized melting aluminum-rich chromium deposits in mantle peridotite (clinophyllite-lenticular dunite) with a high degree of local melting; the second is produced in ultramafic-mafic accumulationsMagma pile crystalline chromium mineralization (spot) at the base of the complex. The former is represented by the 3756 medium-sized chromium deposit. In addition, there are also small deposits such as Jidong, 620, and 820, which have greater economic value. The latter is represented by D2 triangle point mineralization and has little industrial significance. Now, taking the 3756 chromium deposit as an example, we will briefly describe its deposit characteristics.
Hegeola 3756 chromium deposit: 830m long on the surface, 110~300m wide, with an area of about 0.13km2, lying sideways in the northeast direction to a vertical depth of 440m (north of line 4), along the axial direction The maximum extension has been controlled to 930m, and the maximum depth along the slope is 280m. The main ore body group is in the southwest (line 25-27), generally forming a northeast-trending ore-bearing belt. It consists of 186 ore body groups, of which only about 20 are exposed on the surface, and the rest are all blind ore bodies (Figure 3-40, Figure 3-41). Almost all the ore bodies are produced in lenticular dunite in the IV lithofacies zone. The overall axial direction of the ore group is 50° to 60°, tilting to the southeast with an inclination angle of 20° to 70°. The shape of the ore bodies is relatively complex, with most of them being lenses, and the vein-like ore bodies are relatively stable.
The natural types of ores are mainly dense disseminated ores, accounting for 62.04%; medium disseminated ores account for 29.69%; dense massive ores account for 4.1%; sparse disseminated ores account for 3.67%. Ore bodies are produced in lens-like, pod-like, vein-like and cyst-like shapes. The ore body thins rapidly, pinches out and reappears, and its occurrence changes sharply, resulting in a complex ore body morphology.
The ore bodies are mainly made of dunite as the immediate surrounding rock, and a few are clinopyroxene. The contact relationship between the ore body and the surrounding rock shows a clear and rapid gradient or sudden change relationship. The surrounding rocks near the ore are highly altered and are usually dense cryptocrystalline-microcrystalline chlorite aggregates. Most chromium ore bodies are surrounded by serpentine-chlorite thin shells of several centimeters to tens of centimeters, which are similar to Dunite shows a gradual transition relationship.
Figure 3-40 Geological map of Hegeola 3756 chromium deposit
(According to Zhang Guowei et al., 1996)
Ⅴ— Gabbro and altered gabbro; Mg—magnesite; φ2—clinogite; φ1—dunite; 1—section line; 2—chromium ore body and number; 3—chromium ore point; 4 — Occurrence; 5 — Measured faults; 6 — Inferred faults
Figure 3-41 Line B longitudinal section of Hegeola 3756 chromium deposit
(According to the 109 Geological Team of Inner Mongolia Bureau of Geology and Mineral Resources, 1976, abbreviated)
1—Quaternary Series; 2—Paleogene-Neogene Pliocene red clay; 3—Dune; 4—Oclinic gabbrocite; 5-clinobistolite; 6-gabbro; 7-chromium ore body; 8-exploration line number
Chromium spinels are mostly semi-euhedral crystals and other-shaped granular shapes The aggregates are scattered in gangue minerals such as serpentine and chlorite. The particle size is greater than 1mm. Dust-like magnetite is often in the form of particle aggregates, displacing chromium spinel along the edges and cracks of chromium spinel or together with gangue minerals, forming residual structures and network structures. The ore structure includes bean-like or nodular structures, most of whichIt is a dense disseminated ore, and a small part of it is massive ore; disseminated spot structure, dense disseminated and fine-medium-grained disseminated ore have such structure, which is the main structural type of the deposit; strip structure, strip width 2 ~5mm, alternately arranged with dunite, and the strips are chromium spinel aggregates with varying degrees of density; disseminated structures containing massive ores often appear in deep ore bodies.
Ore minerals: The metallic minerals are mainly chromium spinel, followed by dust-like magnetite, and very small amounts of pyrite, chalcopyrite and hematite. The gangue minerals are mainly serpentine, followed by chlorite.
Chromium spinel is mainly magnesia aluminum chromite, followed by chromium-rich spinel and iron-rich chromite. Among different ore types, the chemical composition of chromium spinel changes as follows: from sparse disseminated to dense massive ore, with the ratios of aluminum and magnesium, Mg/(FeO), and A12O3 increasing. The average chemical composition of the ore: Cr2O3: 23.62%, Fe2O3: 1.4%, A12O3: 13.27%, SiO2: 16.52%, FeO: 9.48%, MgO: 25.92%, CaO: 0.57%, K2O: 0.011%, Na2O: 0.076%.
Mineralization conditions: According to research by Bai Wenji (1986), under the low pressure conditions of the seafloor environment, chromium ore bodies in pile crystal rocks are formed at about 800°C. The oxygen fugacity (ƒo2) of chromium ore formation ) is 10-6.
(5) Ore deposit origin and mineralization model
1. Mineralization stages and mineralization eras
There are two genetic subtypes of the Hegenshan chromium deposit Chromium ore: One is a local melting-renovated aluminum-rich chromium deposit (referred to as "local melting type") produced below the Moho surface of the ophiolite section rock. The other is the magma crystallization-segregation type chromium mineralization (spot) produced above the Moho surface of the rock and at the bottom of the ultramafic-mafic accumulation complex (referred to as the "pile crystal type").
Mining age: There are no reports on the metallogenic age of chromium spinel. Generally, the dating data of basic-ultrabasic rocks in the ophiolite set are used to represent the formation age of chromium ore. For " For heap type chromium ore, the diagenesis and mineralization times are relatively close. As for "local melting type" chromium ore, the ore may be formed earlier than the surrounding rock (parent rock). The radiolarian siliceous rock mined from the mining area in this area was identified by Wang Naiwen as Late Devonian (Bao Peisheng et al., 1996). Bao Zhiwei, Chen Senhuang et al. used ultrabasic rock whole rocks in Hegenshan and obtained the Sm-Nd isochron age of 403±27 Ma. The K-Ar ages of the Wusni black clinolite whole rocks were 430 Ma and 285 Ma. (Bao Zhiwei, 1994). Miao Laicheng et al. (2005) obtained three zircon isotope ages in Hegenshan, namely: gabbro 295 Ma, basic dike 298 Ma, and basic lava 293 Ma, which is equivalent to the Late Carboniferous. Therefore, the formation time of the rock belt should be appropriately determined as the Late Devonian to the early Late Carboniferous.
2. Origin of mineral deposits and mineralization factors
The minerals for the two types of chromium ores in the Hegenshan chromium deposit come from both sourcesSpinel lherzolite locally melted deep in the mantle. The formation process can be summarized as follows: ① When the degree of melting of the upper mantle is high, magnesian basic-ultrabasic magma (such as basaltic magma-picrritic magma) can be formed, enter the crust along appropriate channels, and converge into a "subcrustal magma chamber" "Due to differential condensation, the chromium spinel crystallized from the magma sinks to the bottom of the magma chamber due to gravity. As long as the temperature does not drop rapidly, as the crystal accumulation increases, it can eventually form an accumulation with the characteristics of a magma deposit. Layered chromium ore; ② The material remaining in the mantle undergoes a zone-by-zone "fractionation-refining-extraction" process through partial melting in the "convective" movement, and is removed of low-melting-point components, thus forming oblique glow. The "mantle residue" (depleted mantle rock) represented by the peridotite-clinogite-dunite combination. This process can be regarded as a kind of crust-mantle differentiation in the form of "thermal perturbation" or "dynamic perturbation". During this process, the chromium-containing pyroxene and accessory mineral chromium spinel in the mantle rock undergo phase changes to form a chromium-rich slurry. The slurry is initially dispersed as small droplets in the melt and is transported by the plastic flow of the mantle and gravity. Move and gradually become larger, thereby further gathering to form chromium ore, that is, "local melting type" chromium ore. Bao Peisheng et al. (1999) studied mineral deformation and found that the paleo-stress difference and strain rate of the Hegenshan rock mass were both low. It is detrimental to the convergence of chromium ore slurry droplets.
(6) Prospecting model and prediction elements
1) Searching for magnesian ultrabasic rocks is the most important prerequisite for ore prospecting. Faulting activity plays an important role in the emplacement of rock masses. It is not only the basic cause of rock mass emplacement, but it may also destroy the structure of the original crust-mantle transition zone. Attention should be paid to those subduction zones or suture zone structural zones that may mark the main remnants of early oceanic crust, followed by back-arc basin fold zones that have been folded and closed after secondary expansion, especially some fault zones on their edges, which are often conducive to the formation of Mine;
2) By studying the properties of the ophiolite volcanic rocks, we can analyze and judge the location and properties of the original ocean floor structure represented by the "fragments", which will be helpful to find the ancient The "relics" of structural parts with high heat flow activity - favorable areas for mineralization;
3) Rock mass emplacement is a random interception. Therefore, the shape, scale and occurrence of the rock mass are not decisive factors. . The mineralization of rock masses depends on the degree and nature of early "crust-mantle differentiation". Good differentiation has the following two characteristics: first, the well-preserved "undercrustal magma chamber" (with the development of pile crystal rocks), and second, clear lithofacies zoning. On the basis of clarifying the zoning characteristics of lithofacies, restore the original crust-mantle transition zone as much as possible, and then search for new chromium ore bodies.
4) In view of the fact that most of the 3756 deposit is a blind ore body, it has been proved in practice that the magnetic method is more effective in delineating the boundaries of the rock mass. The gravity and torsion scale are only reflected on a few known chromium ore bodies, while Electrical methods and geochemical exploration are not effective. As shown in the comprehensive north-south gravity and magnetic anomaly profile of No. 5 ore body (Figure 342), there are obvious abnormal reflections in gravity and torsion scale, Δg=0.16×10-5m/s2, Vxz=60×10-91/s2, VΔ=116×10-91/s2. The Δg curve is basically symmetrical, and the ore body shows a transition zone with high gravity and magnetic anomalies from low to high.
Figure 3-42 Comprehensive profile of gravity and magnetic anomalies on the 3756 chromium ore body
(According to Liao Changqing et al., 1996)
Q - Quaternary; φδ - dunite; φω - clinopyroxenite
As shown in Figure 3-43, the 620 ore body trends 30° to the northeast and consists of three larger It is composed of a lenticular chromium ore body, with a strike of 40m and an extension of 25m. It tilts southeast with a steep dip angle. The mine roof is buried 4 to 5m deep, and the ore is mainly fine-grained and densely disseminated. It is reflected as gravity anomaly on the ore body, but the magnetic anomaly characteristics are not obvious. However, in areas with complex terrain conditions, the above methods will be subject to certain limitations. Therefore, it is necessary to further explore new methods in order to achieve better results in finding blind ore bodies.
2. Solon Mountain Chromium Deposit in Urad Middle Banner, Inner Mongolia
The Solon Mountain Chromium Deposit is located in the western section of the suture zone between the Siberian Ancient Plate and the Sino-Korean Ancient Plate - Solon Mountain. Lunshan ophiolite belt. Its formation time should be earlier than the Devonian, and its emplacement time is approximately the Late Carboniferous-Early Permian. The rock mass is composed of depleted mantle peridotite and heapite. Chromium ore occurs without exception in dunites of various lithofacies. Among them, the pod-shaped chromium ore formed by local melting of mantle rocks has industrial value. The ore type varies with different rock masses. Generally speaking, it is still a high-aluminum chromium ore, and the scale of the deposit is small.
(1) Overview
Solunshan chromium ore is produced 90km northeast of Urad Middle Banner, Bayannur City, Inner Mongolia. Geographical coordinates: 117°18'00" east longitude , 49°26'00" north latitude. This area produces chromium deposits (points) such as Solon Mountain (Chahanhule, Abu Chromium, and Uzhur). Ore bodies with industrial value are mainly distributed in the Chahanhule mining area in the western section of the Solon Mountain rock mass, the Chahannulu mining area in the middle section, and the Tugemu mining area in the eastern section. Most of them are refractory grade chromium ores, and less are metallurgical grade chromium ores. Cumulative proven C-level reserves are 100,000 tons, and D-level reserves are 392,000 tons, totaling 492,000 tons. As of 1993, C-level reserves were 77,000 tons, and D-level reserves were 383,000 tons, totaling 460,000 tons. They are small chromium deposits.
Figure 3-43 Comprehensive gravity magnetic profile of 620 ore
(Quoted from Liao Changqing et al., 1996)
TCr-Collapse Chromium ore; 1 - Quaternary; 2 - clinobite; 3 - gabbro; 4 - fault; 5 - borehole; 6 - shallow well
The deposit was discovered in 1957 by Inner Mongolia A regional team from the Geological Bureau discovered that ultrabasic rock masses were delineated by aeromagnetic and mantle magnetic surveys in 1958. From 1958 to 1963, the Inner Mongolia Geological Bureau organized general survey and exploration and submitted a reserve report. During this period, some scientific research units in Beijing and Tianjin conducted relevant scientific research work. Sporadic mining began in 1962. In 1982, the Utla Zhong Banner Planning Commission invested in and built a factory.In 1994, 120,000 tons of ore were mined, and the concentrate powder was sent to chemical plants in Tianjin and Hebei. Some of the rich ore was sold to Taiyuan Iron and Steel Co., Ltd.
(2) Regional geological overview
The Solon Mountain ophiolite belt starts from Habutgei in the west, and goes eastward through Solon Aobao, Abugai Aobao (i.e. Abuge), From Uzhur to Har Tolgoi, it is a narrow strip of more than 100 kilometers long from east to west. It has been identified that there are five large ultrabasic rock bodies in the rock belt: Solon Mountain, Abger, and Uzhur. Zhuer, Pingdingshan, and Haye (Figure 3-44) are distributed on the north side of the Solon Aobao-Mandula fault, and are all in fault contact with the surrounding rocks. Most scholars believe that the eastern extension of this rock belt can be connected with the Xilamulun ophiolite belt (Li Jinyi et al., 1987, 1998); Lu Songnian et al. pointed out that the Erlian-Hegenshan ocean basin in the Xingmeng orogenic system to the early mud The Basin Epoch has been closed, and the Ergun, Songliao, Jiamusi and other blocks have been collaged and combined. At this time, only the Xilamulun Ocean Basin remains in the eastern section of the Paleo-Asian Ocean. By the middle of the Carboniferous to the Early Permian, oceanic crust subsidence and continental accretion were coming to an end. The most important geological event was the full-scale collision and docking of Siberia and North China. The Solon Mountains-Xila Mulun Mountains were the Siberian paleoplate and the Sino-Korean paleoplate. The dividing line was also the area where the ancient Asian Ocean finally died. A series of new research results show that the final closure of the ancient Asian Ocean was at the end of the Permian (248±4ma) (Lu Songnian, 2010, internal data). According to the "China Mineralizing Zone Division Plan" (Xu Zhigang et al., 2008), the Solonshan chromium deposit is located in the Bainaimiao-Xilinhot Fe-Cu-Mo-Pb-Zn-Mn-Cr III-level mineralization belt (III-49) (Ⅴm-Ⅰ).
The formation age of the ophiolite belt is Silurian-Devonian. According to Tao Jixiong et al. (2004), the single-grain zircon U-Pb age value measured from the olivine gabbro in the rock mass is (433.6 ±3.6) Ma, intruded in the Carboniferous-Permian.
(3) Rock mass characteristics
The Solon Mountains, Abu Ge (i.e. Abqai Obo) and Uzhuer rock bodies are the larger and larger ophiolite belts in the Solon Mountains. Three structural rock blocks with good mineralization.
The Solon Mountain rock mass is 32km long from east to west, 2 to 6km wide from north to south, and covers an area of 72km2. It enters Mongolia to the north. It is composed of clinolite and dunite, and locally contains a small amount of lherzolite. . Dunite is widely distributed in clinopyroxenite in the form of complex and changeable lens-shaped, strip-shaped, and vein-like bodies, accounting for about 10%. The main body is clinopyroxenite, and the boundary between the two is distinct. The lithofacies zoning of the rock mass is not obvious. Only a small amount of clinopyroxenite-lherzolite is found north of the main peak of Solon Mountain, and the rest are all clinopyroxenite-dunite lithofacies.
The Abger rock mass is located 28km east of the Solon Mountain rock mass. The rock mass is 9.5km long from east to west and 0.18-3.5km wide. It is mainly composed of dunite and clinopyroxenite, with a small amount of isolenite. Peridotite, pyroxenite, gabbro and other lenses. Dunite exposures account for 30% of the total area of the rock mass. From south to north, dunite-clinogite rock facies belts and dunite rock facies belts appear alternately. The surface of the rock mass is highly weathered, mainly silicification and carbonation., the depth of alteration can reach tens of meters.
The Wuzhuer rock mass is located 18km east of the Abuger rock mass, 3.5km long from east to west, and 0.7~1.5km wide. The rock mass is composed of dunite-clinogite, gabbro and pillow Composed of lava, a relatively complete ophiolite combination is exposed in this area, but there is a lack of rock wall groups. It can be divided into three lithofacies belts: dunite lithofacies belt, dunite-clinogite complex lithofacies belt and lherzolite lithofacies belt.
There is not much difference in the chemical composition of the rocks between the three rock masses. In comparison, the chromium content of the Solon Mountain rock mass is slightly higher than that of Abger and Wuzhuer. The serpentinization of the rock bodies is strong, especially the Abger rock mass. In addition to serpentinization, it also has strong silicification and carbonation.
(4) Geological characteristics of ore deposits
Most of the industrial ore bodies in the Solon Mountain rock mass are, without exception, produced in the clinolite-dunite complex belt. Among dunite lenses, they appear in groups and are distributed in bands, with most of them in the north and decreasing toward the south. From west to east, it includes three mining areas: Chahanhule, Chahannulu and Tukemu (Figure 3-44, Figure 3-45). A total of more than 100 ore bodies have been discovered, of which nearly 80% are blind ore bodies. Ore body shapes are complex and diverse, mainly lenticular, pod-shaped, vein-shaped, reticular-vein-shaped, cyst-shaped, nest-shaped, cylindrical and irregular. The scale can range from tens to 300m. The ore type is mainly disseminated, and the ore-making chromium spinel is chromium ore (chromium-rich type). Cr2O3 varies greatly, ranging from 10.68% to 31.59%, and can be greater than 40% to 50% in some cases; Cr2O3/＜FeO＞ can It reaches 3.03~3.06. The proven reserves account for 73.45% of the entire area (the entire rock mass of Solon Mountain).
Figure 3-44 Geological Map of Chahanhule Mining Area of Solon Mountain Rock Mass
(According to Geological Research Team of Inner Mongolia Bureau of Geology and Mineral Resources, 1984)
p>1—Clinolite; 2—Dune; 3—Chromium ore body; 4—Fault; 5—Ore body occurrence; 6—Exploration profile line
Abg Rock mass mineralization is relatively common. There are two main ore groups, Cr209 and Cr207. A total of 37 ore bodies have been discovered, and only 4 are exposed on the surface. The ore bodies are generally tens to more than a hundred meters long and several meters thick. The proven reserves account for 9.8% of the entire region.
Wuzhuer rock body has two ore groups I and III, consisting of 18 ore bodies. Only 2 ore bodies are exposed on the surface, and the others are blind ore bodies. The proven reserves account for 16.9% of the entire region.
The mineral composition of the ore is mainly aluminum chromite, which contains slightly higher iron than the Hegenshan chromium ore, which may be related to the stronger alteration effect. In addition, it contains small amounts of magnetite, hematite, pyrrhotite and pentlandite. The gangue minerals are mainly serpentine, followed by chlorite and magnesite.
The ore has a heteromorphic-semi-euhedral crystal structure, a fragmented structure, a chipped structure, a metasomatic structure, a network-vein structure and a plastic deformation structure. The ore structure is complex, mainly disseminated ore, followed by dense massive, striped, nested, cystic, mottled and reverse mottled structures.
Figure 3-45 Chahannulu Mining Area in Solon MountainⅡ Ore Group 4 Exploration Line Profile
(According to the Geological Research Team of Inner Mongolia Bureau of Geology and Mineral Resources, 1984)
1—Clinolite; 2—Monolithite Rock; 3—carbonated ultrabasic rock; 4—diabase; 5—chromium ore body; 6—inferred fault; 7—lithofacies boundary; 8—boring hole
(5) Prospecting suggestions
Inner Mongolia is one of the earliest regions in my country to carry out chromium ore prospecting work, and one medium-sized mineral deposit; two small-scale mineral deposits; mineral points, There are 36 mineralization points, nearly 80% of which are blind ore bodies. It is enough to prove that this area has good prospecting potential. The following suggestions for prospecting are put forward:
1) The chromium ores with industrial value in Inner Mongolia are all ophiolite-type chromium ores. The Solon Mountain ophiolite belt is also the suture zone between the Siberian paleoplate and the Sino-Korean paleoplate. Therefore, we should strengthen the research on regional ophiolite belts, search for favorable areas for mineralization, and increase the intensity of prospecting and exploration, hoping to achieve new progress.
2) Strengthen method research. The success of prospecting methods can often lead to breakthroughs in prospecting. In view of the large area covered by the Quaternary System in Inner Mongolia, most of the ore bodies found in the past are blind ore bodies, and a certain amount of experience has been accumulated. It is necessary to further strengthen the application of comprehensive geophysical methods in the search for chromium ores, and strive to make some achievements. breakthrough.
3) Pay attention to new mineralization theories at home and abroad, conduct in-depth research on the mineralization rules of China's chromium ore, and combine it with the trend of deep mineralization, focusing on the technical methods of deep mineralization, and lead from point to area.
『五』Types of mineral resource/reserve boundary lines
Mineral resource/reserve boundary lines have different names because of their different meanings. The mineral resource/reserve boundaries are generally delineated first. Determine the mineral resource/reserve boundary (or boundary base point) in a single project, and then connect the corresponding points on the profile or plane to form a boundary based on the resource/reserve boundary base points in all projects to obtain various ore bodies in the three-dimensional space. boundary.
1. Zero point boundary line
The zero point boundary line is a line connecting points on the projection plane where the thickness of the ore body or the content of useful components tends to zero. That is, the line connecting the pinch out points of the ore body. The zero-point boundary line is often used as an auxiliary line to determine the recoverable boundary line, rather than the true resource/reserve boundary, because it is impossible to calculate mineral reserves to the zero-point boundary.
2. The recoverable boundary line
The recoverable boundary line is a line connecting the base points determined by the minimum recoverable thickness and the lowest industrial grade or the lowest industrial meter percentage. It is used to delineate the boundary location of industrial ore bodies, that is, the ore within the mineable boundary is regarded as reserves or basic reserves.
3. Ore grade and type boundary lines
Within the scope of the mineable boundary line, different technical grades and ore types are divided according to the requirements of the ore technical grade and type. dividing line. Indicates the distribution of various grades and types of ores in industrial ore bodies.
When determining the boundaries of ore grade and natural type, it must be within the range of the mineable boundary.Pay attention to the geological factors that control ore grade and natural type. For example, when determining the boundary between the oxidized zone and the primary zone, it must be considered that the boundary between the oxidized zone and the primary zone is mainly controlled by the groundwater level, and the groundwater surface can be regarded as horizontal within a short distance, so its boundary should be horizontal.
Table 8-7 Standard table for the classification of natural types of general non-ferrous metal ores
The classification of primary ores and oxidized ores of iron ores is generally based on TFe/ Measured by different ratios of FeO. When the iron-containing minerals in iron ore deposits are mainly magnetite and are later oxidized into hematite and limonite:
Primary ore: TFe/FeO ＜ 2. 7; mixed ore: TFe /FeO2. 7 ~ 3. 5; Oxidized ore: TFe/FeO> 3. 5.
When the iron-containing minerals in the ore are mainly siderite or magnetite ore with a high iron silicate ratio, the classification standards for primary ore and oxidized ore will be considered separately.
4. Reserve category boundary lines
That is, the boundaries demarcated according to the conditions of different reserve categories, such as the dividing line between reserves, basic reserves and resources.
5. Inner boundary line and outer boundary line
The inner boundary line is the boundary connecting the control points of the ore encountering project on the edge of the ore body. It represents the part of the ore body controlled by the exploration project. The distribution range of the ore body; the outer boundary line is a boundary line determined outward or in depth based on the edge ore mining project to indicate the possible distribution range of the ore body. Spatially speaking, the zero point boundary line belongs to the outer boundary line, while other types of boundary lines can be within the inner boundary line or between the inner and outer boundary lines.
6. Temporarily unminable boundary line
This boundary line is delineated based on the boundary grade. The amount of mineral between this line and the mineable boundary line is the resource amount.
『Continent』 Material composition of igneous rock
The material composition of igneous rock is the most basic characteristic of igneous rock. It is not only the basic basis for the classification and naming of igneous rock, but also the basis for studying the origin, generation and evolution of magma. important means.
1. Chemical composition of igneous rocks
Research shows that all elements in the earth's crust appear in igneous rocks, but the contents vary greatly. According to the content and geochemical significance of elements in igneous rocks, they are divided into major elements, trace elements and isotopes.
(1) Main elements
There are many kinds of elements that make up igneous rocks, but O, Si, Al, Fe, Mn, Mg, Ca, Na, K, Ti, P 12 elements including H and H are the main ones, and the O element has the highest content, reaching more than 45%. The sum of these 12 elements accounts for more than 99% of the total mass of igneous rocks and is called the main rock-forming elements. When studying igneous rocks, their chemical composition is not expressed in the form of elements, but in the form of oxides, namely SiO2, TiO2, Al2O3, Fe2O3, FeO, MnO, MgO, CaO, Na2O, K2O, P2O5, H2O. These oxygenThe content of chemical compounds in igneous rocks is usually greater than 0.1% and is called the main rock-forming oxide (Table 2-1). According to research needs, CO2, Cr2O3 and other contents can also be given.
Table 2-1 Chemical composition of igneous rocks in China (wB/%)
(1) SiO2 has the largest content and the largest variation range among igneous rocks. The most important oxide. According to SiO2 content, igneous rocks are divided into four types: acidic rock (SiO2>63%), neutral rock (SiO252%~63%), basic rock (SiO245%~52%) and ultrabasic rock (SiO2<45%) type. The acidity or basicity of igneous rocks is usually referred to by the SiO2 content. The higher the content, the greater the acidity and the lower the basicity of the rock. Research on the chemical composition of igneous rocks shows that as the SiO2 content changes, the contents of other major rock-forming oxides change regularly (Figure 2-1). As the SiO2 content increases, the Na2O and K2O contents gradually increase, and the FeO and MgO contents continue to decrease; while the CaO and Al2O3 contents increase rapidly from ultrabasic rock to basic rock as the SiO2 content increases, and then from basic rock to neutral rock. , it gradually decreases when the acidic rock changes. In addition to silicate minerals such as feldspar, mica, amphibole, and pyroxene, SiO2 in magma appears as an independent quartz mineral when there is excess.
Figure 2-1 The relationship between SiO2 content and other oxides in igneous rocks (Qiu Jiaxiang, 1985)
(2) The relationship between Na2O and K2O content and is called the total alkali content, which has great differences in different lithologies (Table 2-1). Na2O and K2O are the main components of alkaline feldspar. When the total alkali content is high, alkaline dark minerals and parafeldspar can appear in the rock. In the study of igneous rocks, the Rittmann index (σ) is commonly used to classify the alkaline degree of rocks, σ=[w(Na2O+K2O)2]/[w(SiO2)-43%], and rocks with σ<3.3 are calc-alkali. For alkaline rocks, rocks with σ=3.3~9 are alkaline rocks, and rocks with σ>9 are peralkaline rocks. However, for rocks with high SiO2 content (SiO2>70%), the Rittmann index is ineffective in determining whether it is alkaline or sub-alkaline (Deng Jinfu et al., 2004). This is because the dilution effect of SiO2 will cause the alkali content to be relatively biased. If it is low, the calculated σ value is too small, and it will be mistakenly determined as a calc-alkaline rock series. For example, some alkaline rhyolites with SiO2>80% should be paid special attention to.
(3) Al2O3 is the rock-forming oxide second only to SiO2. The content of Al2O3 in igneous rocks is mainly between 10% and 18%. Al2O3 combines with SiO2 and CaO, Na2O, K2O to form plagioclase, alkali feldspar and feldspar-like minerals; it combines with FeO, MgO, CaO and SiO2 to form minerals such as pyroxene, hornblende and biotite. Al2O3 also plays an important role in the classification and origin research of igneous rocks: ① According to the alkali content and CThe relative ratio between aO and Al2O3 content divides igneous rocks into peralkaline rocks (Al2O3
(4) MgO, FeO and SiO2 combine to form iron-magnesium silicate minerals, such as olivine, pyroxene, etc. Because MgO, FeO and SiO2 content are negatively correlated (Figure 2-1), olivine and pyroxene appear only when the SiO2 content is low. The classification of igneous rock series, rock classification and genesis research based on the main elements is one of the main methods of igneous rock research and involves a lot of content. For beginners, the following three applications related to major elements should be mastered and understood.
1. Division of igneous rock series
Igneous rocks can be divided into three rock series, namely alkaline, calc-alkaline and tholeiitic basalt series, the latter two together are called Sub-alkaline series. First, according to the silica-alkali diagram (Figure 2-2), distinguish the alkaline series (A) and the sub-alkaline series (S). For sub-alkaline series rocks, use the TFeO/MgO-SiO2 diagram and the TFeO/MgO-TFeO diagram (Figure 2-3) or the AFM diagram (Figure 2-4) to further distinguish whether it is the tholeiitic basalt series or the calc-alkaline series. Subalkaline series igneous rocks can also be divided into low-potassium, medium-potassium, high-potassium and potash rock types based on the SiO2-K2O diagram (Figure 2-5).
Figure 2-2 Silica-alkali diagram (Irvine, 1977)
Figure 2-3 TFeO/MgO-SiO2 diagram of volcanic rock series division (a), TFeO/MgO-TFeO diagram (b) (Miyashiro, 1974)
Figure 2-4 AFM diagram of the division of igneous rock series (Rollison, 1993; quoted from Yang Xueming et al., 2000)
Figure 2-5 SiO2-K2O diagram of subalkaline igneous rock series division (LeMaitree et al., 1989; Rickwood, 1989)
There are many other diagrams related to the classification of igneous rock series. When applying, special attention should be paid to the conditions of use of each diagram and cannot be copied mechanically. For example: When applying the silicon-alkali diagram to divide rock series, you need to be careful with high-silica granite and rhyolite (generally SiO2>70%). Because the high SiO2 content leads to low alkali content, it is the same as the Rittmann index. When determining Alkaline or sub-alkalineIt will be invalid in the alkaline series (Deng Jinfu et al., 2004), causing alkaline granite and alkaline rhyolite to fall into the sub-alkaline series area, which is obviously wrong. In the original diagram, the upper endpoint of the dividing line between the two series ends at the SiO2 content of 67%, and it does not extend upward for this reason. Irvine & Baragar (1971) gave the mathematical equation of the graphical dividing line as: S=-(3.3539×10-4)A6+(1.2030×10-2)A5-0.15188A4+0.86096A3-2.1111A2+3.9492A+39. In the formula, S=w(SiO2), A=w(Na2O+K2O), when the SiO2 in the rock is greater than S calculated by the formula, it is a sub-alkaline zone, and vice versa, it is an alkaline zone. Deng Jinfu et al. (2004) suggested that the alkalinity rate proposed by Wright (1969) [AR=w(Al2O3+CaO+Na2O+K2O)/w(Al2O3+CaO-Na2O-K2O) should be used to classify igneous rock series with SiO2>70% ], and use the SiO2-AR diagram to distinguish them (Figure 2-6). At the same time, the appearance of alkaline dark minerals in rocks is the most important petrographic sign, and the over-alkali index ([(Na2O+K2O)/Al2O3]>1, number of molecules) is the most reliable way to identify alkaline granite (rhyolite) Geochemical parameters.
Figure 2-6 Illustration of the SiO2-AR relationship of the igneous rock classification method (Wright, 1969; quoted from Deng Jinfu et al., 2004)
Research shows that, Different series of igneous rocks have many differences in their magma origin, evolution, and structural background. Therefore, accurate rock series classification is helpful in determining the origin of igneous rocks. The relevant content will be introduced in subsequent chapters.
2. Harker type rock chemical composition variation diagram
This is the simplest but often used diagram. This diagram often uses SiO2 or MgO content as the abscissa, and other The main oxide content is a diagram composed of the ordinate (Figure 2-7). According to research needs, relevant parameters, such as alkalinity rate (AR), differentiation index (DI=Q+Or+Ab+Ne+Lc+Kp, standard mineral), etc., can also be selected as variables to study. Generally, the oxide data used should be the oxide content recalculated after removing H2O, loss on ignition, etc. from the complete silicate analysis. This diagram shows the changing trend of other oxides or parameters as the SiO2 or MgO content changes (Figure 2-7). Generally, if there is a strong linear correlation between the chemical compositions of igneous rocks that coexist closely in the same area, spatially, and have large composition changes on the Harker diagram, it indicates that these rocks are likely to be a group of rocks formed by the evolution of homologous magma. . If there is no correlation, it means that they may be the product of different magma crystallizations.
3. Calculation and application of CIPW standard minerals
When igneous rocks rapidly condenseWhen formed under conditions, the crystalline mineral particles are small, or are partially or even entirely composed of glass (such as many volcanic rocks), then the actual mineral composition and content of the rock cannot be known, and rock classification and naming rely on the actual mineral composition and content. invalid. In order to solve this problem, people have proposed a method of using chemical composition to calculate the ideal mineral composition and content in igneous rocks, that is, the standard mineral calculation method. At present, the widely used calculation method is the method jointly proposed by Cross, Iddings, Pirsson, and Washington (1902) in the United States, referred to as the CIPW standard mineral calculation method.
This method is based on the results of experimental research on the crystallization sequence of minerals in anhydrous magma, and formulates standard minerals according to ideal molecular formulas. First, the oxide mass percentage of the rock is converted into the number of oxide molecules. Then, in a certain order, the number of molecules is combined into several standard mineral molecules with ideal composition according to certain rules. Finally, the number of standard mineral molecules is converted into the standard Mineral content percentage. For detailed calculation procedures, see "Igneous Rock Petrology" edited by Qiu Jiaxiang (1985) and "Igneous Rock Petrology and Petrology" edited by Lin Jingqian (1995). Now, relevant software has been compiled and can be completed quickly through computers. CIPW standard minerals can roughly reflect the mineral composition of rocks, but they are not necessarily the minerals that actually appear in rocks. The calculation results are used in rock classification (Figure 2-8, Figure 2-9), magma formation or crystallization temperature and pressure conditions determination (Figure 2-10) and many other aspects.
Figure 2-7 Harker diagram of volcanic rocks in Mazama Mountain, Oregon, USA (Winter, 2001)
Figure 2-8 Standard mineral classification of basalt Illustration (Qiu Jiaxiang, 1988)
Figure 2-9 An-Ab-Or classification diagram of the standard minerals of granitic rocks (Rollison, 1993; quoted from Yang Xueming et al., 2000 )
Figure 2-10 Determination of source depth of acidic magma (Winter, 2001)
(2) Trace elements
Trace elements refer to those elements whose content is very small in rocks. Their content can only be expressed in parts per million (10-6). Generally, their total amount is <1%. The study of trace elements has become a key component of modern petrology and can distinguish the petrogenetic and evolutionary processes more effectively than major elements. Frequently mentioned trace elements are vanadium (V), cobalt (Co), nickel (Ni), chromium (Cr), rubidium (Rb), strontium (Sr), barium (Ba), cesium (Cs), thorium ( Th), uranium (U), zirconium (Zr), hafnium (Hf), niobium (Nb), tantalum (Ta) and rare earth elements (REE), etc. Trace elements usually do not appear as independent minerals. They mainly replace the main elements in minerals in the form of homogeneous isomorphs, such as Cr and Ni.On behalf of the Mg and Fe positions in olivine and pyroxene, Sr can occupy the Ca position in plagioclase, etc.; secondly, it exists in rapidly condensing volcanic glass and gas-liquid inclusions; thirdly, it is adsorbed on the mineral surface.
Trace elements in igneous rocks often change regularly with changes in the content of major rock-forming elements. For example, as the acidity of the rock increases, the content of siderophile elements (V, Cr, Co, Ni, etc.) decreases, while the alkali metal trace elements (Li, Rb, Cs) increase. Studying the characteristics of trace elements can obtain important information about the division, origin and evolution of rock series.
Rare earth elements include lanthanide elements with atomic numbers 57 to 71: lanthanum (La), cerium (Ce), praseodymium (Pr), neodymium (Nd), promethium (Pm), samarium (Sm) , europium (Eu), gadolinium (Gd), terbium (Tb), dysprosium (Dy), holmium (Ho), bait (Er), thulium (Tm), ytterbium (Yb), lutetium (Lu), in addition, usually Count the element yttrium (Y) with atomic number 39 as a rare earth element. Except for Pm, which is an artificial radioactive product, the others are elements with similar geochemical properties, are refractory and symbiotic, and are not easy to migrate in secondary processes. The total amount of rare earth elements, curve partitioning patterns, and europium (Eu) anomalies all contain important information on the origin and evolution of magma and rock formation mechanisms.
During the process of magma crystallization, some trace elements preferentially enter the crystalline mineral phase, or when the source rock is partially melted to form magma, they tend to remain in the source rock mineral. These elements are called phases. On the contrary, during the magma crystallization process, they are not captured or accommodated by the minerals that crystallized early, but are enriched in the residual melt, or when the source area rocks are partially melted to form magma, they preferentially enter the melt phase. , these elements are called incompatible elements, also called hygrophilic elements. It is worth noting that the degree of elemental compatibility and incompatibility varies among different magmas or minerals. For example, P is an incompatible element in mantle magma, but in crustal granite magma, it is a compatible element even if it appears in the form of trace elements; for example, Cr, Ni, and Co elements are compatible with olivine. element, and for plagioclase, it is an incompatible element.
Incompatible elements are further divided into high field strength elements (HFSE) and low field strength elements (LFSE) based on the field strength (charge/radius ratio, i.e. ion potential). Elements with an ion potential greater than 2.0 are called high-field-strength elements, including lanthanide elements, Sc, Th, U, Pb, Zr, Hf, Ti, Nb, Ta, etc.; elements with an ion potential less than 2.0 are called low-field-strength elements. Also called large ion lithophile elements (LILE), including Cs, Rb, K, Ba, Sr, divalent Eu, Pb, etc.
It is precisely due to the above-mentioned petrogeochemical differences in trace elements that there is a strong differentiation phenomenon in the longitudinal section of the lithosphere. For example, the crust formed through magma has much higher abundances of incompatible elements than the mantle. Local interaction of fluids with the mantleThe generation process can cause the enrichment of incompatible elements in the mantle, resulting in the inhomogeneity of the mantle composition. Magma originating from different source areas will inevitably retain traces of trace elements in the source area. Therefore, by studying the trace element characteristics of igneous rocks, we can reveal petrogenetic information such as the nature of the magma source area and the evolution of magma. In the study of the genesis of igneous rocks, trace element ratios and their diagrams are often used for tracing, as well as trace element spider diagrams (Figure 2-11) and rare earth element distribution pattern diagrams (Figure 2-12). Figure 2-11 is a standardized spider diagram of trace elements of mid-ocean ridge basalt (depleted mantle source area), alkaline ocean island basalt (enriched mantle source area), and island arc calc-alkali basalt (fluid metasomatized mantle source area). The difference is obvious. Island arc calc-alkali basalts are depleted of high field strength incompatible elements, especially Nb and Ta; alkaline ocean island basalts are strongly enriched in Nb and Ta; mid-ocean ridge basalts are depleted of large ion lithophile elements Ba, Rb, K. Explain the differences in material composition in the source areas of these three types of basalt magma. Figure 2-12 is the distribution curve of rare earth elements in igneous rocks from different sources (Xu Xisheng and Qiu Jiansheng, 2010). Based on the research results of other scholars, Xu Xisheng and Qiu Jiansheng (2010) concluded that bosonite is The island arc area is formed by the magma crystallization produced by the direct melting of the depleted mantle after being metasomatized by fluids released from the subducted oceanic crust. It has not undergone significant evolution. The total amount of rare earth elements is low, there is no obvious fractionation of light and heavy rare earth elements, and the heavy rare earth elements are slightly enriched. U-shaped curve characteristics; adakite (i.e. island arc dacite) is formed by direct melting of subducted oceanic crust (and its sediments). Light and heavy rare earth elements are strongly fractionated. The rare earth element curve slopes steeply to the right, and heavy rare earth elements are strongly fractionated. Depletion; if this melted ocean crust reacts with mantle peridotite, it can evolve into high magnesium andesite. The "common arc andesite" in Figure 2-12 is formed by crystallization differentiation of basaltic magma and has obvious negative Eu anomalies. For detailed explanations of trace elements, please refer to the textbook "Petrology of Trace Elements in Igneous Rocks" (Li Changnian, 1992) and the book "Rock Geochemistry" (Rollison, 1993; translated by Yang Xueming et al., 2000).
Figure 2-11 Spider diagram of trace elements in basalt in different tectonic environments (Blattetal., 2006)
Figure 2-12 Different origins Rare earth element distribution pattern diagram of types of igneous rocks
(3) Isotopes
Isotopes have been widely used in the study of igneous rocks. They can not only determine the formation age of igneous rocks, but also It can trace the properties of magma source areas and the formation and evolution process of igneous rocks, and explore important scientific issues such as the interaction between crust and mantle and the growth of continental crust. Isotopes can be divided into two categories: stable isotopes and radioactive isotopes.
◎Stable isotopes: The most commonly used isotopes in igneous rocks are O, H and S. Their study can obtain information on the origin of igneous rocks and the origin of magma. The commonly used data include the δ18O value of oxygen isotopes, The δD value of hydrogen isotopes and the sulfur isotopeδ34S value. Taking oxygen as an example, oxygen isotopes are composed of 16O, 17O, and 18O. During geological processes and magmatism, 16O and 18O fractionate due to large mass differences, resulting in differences in the compositions of 16O and 18O at different locations in the lithosphere. The composition of oxygen isotopes is usually represented by δ18O(‰). δ18O=1000×[(18O/16O) sample-(18O/16O) standard]/(18O/16O) standard. The (18O/16O) standard value is usually the average of sea water. value. Although there is minor inhomogeneity in the oxygen isotope composition of the mantle, the δ18O value is basically around 5.7‰±0.3‰. Igneous rocks of different origins have different oxygen isotopes. For example, granite formed by the melting of metasedimentary rocks has δ18O＞10‰; granite formed by the differentiation of mantle-derived magma has δ18O＜6‰.
◎Radioactive isotopes: Radioactive isotopes in igneous rocks include K-Ar, U-Pb, Rb-Sr, 40Ar-39Ar, Sm-Nd, Re-Os, Lu-Hf, etc., and are mainly used to determine igneous rocks. Formation age, origin of tracer rocks, and formation and evolution of the earth's crust. The most commonly used age calculation methods in isotope geological dating are isochron age, model age, U-Pb consistent line age and inconsistent line age. Relevant content is introduced in monographs and textbooks. Please refer to "Chen Yuelong et al. (2005)" Isotope geochronology and geochemistry.
In terms of isotope tracing of the genesis of igneous rocks, commonly used data include the initial ratio of Sr isotopes (87Sr/86Sr)i, 87Sr/86Sr, 143Nd/144Nd, εNd(t), 187Os/186Os, 206Pb/ 204Pb, 208Pb/206Pb, 177Hf/176Hf, εHf(t), etc. The reason why the isotope ratios of igneous rocks can trace the characteristics of the source area is because the mass difference between these commonly used isotope pairs is too small, making it impossible for these isotope pairs to fractionate under the control of the crystal-liquid equilibrium process. They are subsequently differentiated. remain constant during the action. Therefore, the magma formed by partial melting has the characteristics of the isotopic composition of the source area. This fact gave rise to two major developments in isotope geochemistry. One was that specific source areas could be identified by their characteristic isotopic compositions. Figures 2-13 and 2-14 show the Pb, Sr, and Nd isotope compositions of different source areas such as depleted mantle, primitive mantle, enriched mantle, upper crust, and lower crust. The differences are very obvious. For example, granite formed by partial melting of mantle-derived igneous rocks (such as gabbro) has (87Sr/86Sr)i<0.707; granite formed by partial melting of crust-derived argillaceous metamorphic rocks (such as mica schist, aluminum-rich gneiss) , whose (87Sr/86Sr)i>0.708. Second, the mixing and contamination of source intervals with different isotope compositions can be identified. For example, Figure 2-15 is the initial P of Paleogene-Neogene volcanic rocks on the Isle of Skye, Scotland.b isotope composition. The acidic granite and basic volcanic rocks in this area are arranged linearly in the diagram, and are located between the granulite phase lower crust and the Hebrides (ocean island) mantle Pb isotope composition, but deviate from the upper crust. (Figure 2-15a). Therefore, this set of volcanic lava is interpreted to be formed by magma derived from the mantle that was contaminated by the granulite phase of the lower crust.
Figure 2-13 Pb isotope composition in different source areas (Rollison, 1993; cited from Yang Xueming et al., 2000)
Figure 2 -14 Sr-Nd isotope compositions in different mantle source areas (Winter, 2001)
Figure 2-15 Initial Pb isotope ratios of Paleogene-Neogene volcanic rocks on the Isle of Skye, Scotland Illustration (Thompson, 1982)
2. Mineral composition of igneous rocks
(1) Classification of minerals in igneous rocks
Mineral compositions of igneous rocks It not only reflects the chemical composition of rocks, but also represents the temperature, pressure and fluid conditions of rock formation. It is not only the main basis for rock classification and naming, but also an important symbol for judging rock formation conditions. There are many types of minerals found in igneous rocks, but there are only more than 20 common minerals. Among them, the most important minerals that play an important role in rock classification are: olivine family, pyroxene family, hornblende family, mica family, Alkaline feldspar, plagioclase, quartz and feldspar-like (nepheline, leucite), etc., these minerals are called the main rock-forming minerals. In the process of studying igneous rocks, people classify minerals based on their chemical composition, color, content, origin, and role in classification and naming. There are mainly the following classification methods:
1. Composition classification of minerals
According to the chemical composition of the minerals, the minerals in igneous rocks are divided into iron-magnesium minerals and silicon-aluminum minerals.
◎ Ferromagnetic minerals: Minerals with high MgO and FeO contents mainly include olivine (forsterite, chrysolite and fayalite), orthopyroxene (perilla pyroxene, bronze Pyroxene, enstatite), clinopyroxene (ordinary pyroxene, diopside, labile pyroxene and titanium-rich pyroxene), amphibole (mainly ordinary amphibole), biotite, etc. They appear in dark colors such as black, black-green, and dark brown in rocks, so they are also called dark minerals. Dark minerals rich in Na2O are called alkaline dark minerals, such as nite, nite, soda amphibole and soda amphibole.
◎Silicone minerals: minerals that do not contain MgO and FeO and are rich in SiO2 and Al2O3, mainly quartz, plagioclase, alkali feldspar and feldspar. They appear colorless, gray-white and other light-colored tones in rocks, so they are also called light-colored minerals.
The volume percentage of dark minerals in igneous rocks is called color ratio, which is one of the important indicators for the classification and identification of igneous rocks. Igneous rocks with a color ratio of >90 are ultramafic rocks, basic rocks have a color ratio of 40 to 90, neutral rocks have a color ratio of 15 to 40, and acidic rocks have a color ratio of <15.
2. Classification of mineral content and functions
According to the content of minerals in igneous rocks and their role in rock classification and naming, minerals in igneous rocks are divided into main minerals, secondary minerals and accessory minerals. mineral.
◎Main minerals: Minerals with high content in rocks and playing a major role in the classification of rock types. For example, quartz, alkali feldspar, and plagioclase in granite are the main minerals; pyroxene and plagioclase are the main minerals in gabbro.
◎Minor minerals: Minerals that are less abundant in rocks than major minerals and do not play a major role in classifying rock categories, but play a decisive role in determining rock species. For example: a small amount of quartz can appear in gabbro. The presence or absence of quartz does not affect the naming of gabbro, but it does control whether it is called quartz gabbro or quartz-containing gabbro. So quartz is a minor mineral in gabbro.
◎Accessory minerals: The content in rocks is usually <1% and does not affect the classification and naming of rocks. Common ones include magnetite, ilmenite, sphene, zircon, apatite, allanite, monazite, etc.
3. Classification of the origin of minerals
According to the origin of the minerals in igneous rocks, they are divided into primary minerals, diagenetic minerals and secondary minerals.
◎Primary minerals: Minerals formed during the condensation and crystallization process of magma. Most minerals in igneous rocks belong to this category. Primary minerals can be further divided into high-temperature minerals and low-temperature minerals according to their generation environment. Generally speaking, due to the high temperature of volcanic rock magma, the minerals formed are high-temperature types, such as high-temperature plagioclase, high-temperature quartz (β-quartz) and high-temperature alkaline feldspar (lucite); they appear in plutonic intrusive rocks. Low-temperature minerals, such as low-temperature plagioclase, low-temperature quartz (α-quartz) and low-temperature alkaline feldspar (orthoclase).
◎Diagenetic minerals: After the crystallization of magma, due to the continuous decrease of temperature and pressure, the primary minerals are transformed to form new minerals, which are called diagenetic minerals. For example, high-temperature β-quartz transforms into low-temperature α-quartz; high-temperature feldspar transforms into low-temperature orthoclase; orthoclase decomposes to form new striated feldspar; among them, α-quartz, orthoclase and striated feldspar Feldspar is a rock-forming mineral.
◎Secondary minerals: They are post-magmatic minerals. They are new minerals formed after the magma diagenesis due to the replacement and filling of residual volatile matter and post-magmatic hydrothermal fluids. Secondary minerals It is mainly new minerals formed by fluid metasomatism of primary minerals and diagenetic minerals, or new minerals filled in intergranular spaces and pores. ① New minerals formed by metasomatism of primary minerals and diagenetic minerals are also called alteration minerals, which are mainly based on hydration and carbonation. For example, plagioclase undergoes metasomatism to form albite, calcite and zoisite; clinopyroxene is altered into actinolite and tremolite; biotite is transformed into chlorite. ② Secondary minerals filled in pores or gaps, such as zeolite, quartz clusters, etc. filled in pores of volcanic rocks. Secondary minerals also include gas-formed minerals such as fluorite and tourmaline after the magmatic period.
Some alteration and metasomatism are often accompanied by mineralization.Therefore, studying the alteration and metasomatism process is of great significance to the general prospecting of ore deposits after the magmatic period.
(2) The relationship between the chemical composition of igneous rocks and mineral symbiotic combinations
Different types of igneous rocks have different mineral compositions, and different rock-forming minerals form regular symbiotic combinations. On the one hand, its combination is related to the temperature and pressure at the time when the rock was formed; on the other hand, it mainly depends on the chemical composition of the rock. Among the chemical components, the contents of SiO2, K2O+Na2O, and Al2O3 have the greatest impact.
1. The influence of SiO2 content on mineral symbiotic combinations
As mentioned before, SiO2 is the most abundant oxide in igneous rocks. It can form various types of silicic acid when combined with other oxides. Salt mineral. When the SiO2 content is excessive (supersaturated), it will be freed from the silicate melt and crystallize into quartz. Therefore, the appearance of quartz is a sign of SiO2 supersaturation in igneous rocks. When the SiO2 content is insufficient (unsaturated), SiO2 unsaturated minerals appear in the rock, and no quartz is formed, because after these minerals are formed, if there is excess SiO2 in the magma, the two will react to form other minerals, such as:
Petrology
People are accustomed to calling silicate minerals that can coexist with quartz in igneous rocks as SiO2 saturated minerals (or silicic acid saturated minerals) , such as pyroxene, hornblende, plagioclase, alkaline feldspar, mica, etc.; silicate minerals that are not symbiotic with quartz are called SiO2 unsaturated minerals (or silicate unsaturated minerals), such as forsterite , feldspar-like (nepheline, leucite), yellow feldspar, black garnet, etc.; quartz is called a silicic acid supersaturated mineral.
As mentioned in the previous section, the major oxides in igneous rocks show regular changes with changes in SiO2 content. Reflected in the mineral composition, as the SiO2 content increases, the iron and magnesium minerals in the rocks change from more to less, and the mineral types change from olivine and pyroxene to hornblende and biotite; silicon and aluminum minerals appear from scratch, or from From less to more, mineral types evolve from Ca-rich to Na, K, and Si-rich (Figure 2-16).
Figure 2-16 Changes in igneous rock mineral combinations (Adams, 1956)
2. Effect of alkali content on mineral symbiotic combinations
The mineral combinations in igneous rocks with different alkaline (K2O+Na2O) contents are also very different. As mentioned before, according to the size of Rittman index σ, igneous rocks can be divided into calc-alkaline rocks, alkaline rocks and per-alkaline rock types. The mineral combinations of different types of rocks are obviously different. In calc-alkaline rocks with σ<3.3, feldspar, black garnet and alkaline dark minerals (neonite, sodalite, astellite, etc.) do not appear, but feldspar and quartz appear. And ordinary pyroxene, diopside, clinopyroxene and ordinary hornblende, etc. In over-alkaline rocks with σ>9, feldspar-like and alkaline dark minerals (neonite, neonite, amphibole, star leaf, iron-rich mica, etc.) often appear. The feldspar is mainly alkaline feldspar. Black garnet is also more common, and orthopyroxene and quartz are not seen. σ=Among the alkaline rocks of 3.3 to 9, alkaline feldspar and alkaline dark minerals are common, and quartz, feldspar-like (the two are not symbiotic) and plagioclase other than albite can appear.
3. The influence of Al2O3 content on mineral combinations
According to the relationship between Al2O3 and Na2O+K2O, CaO content, igneous rocks can be divided into peralkaline, peraluminous and partial aluminous. There are three types of rocks, and different types of rocks have their own characteristic mineral combinations. Alkaline feldspar, feldspar-like and alkaline dark minerals appear in peralkaline rocks; in addition to feldspar, quartz and biotite, muscovite, topaz, tourmaline, manganese-aluminum garnet appear in peraluminous rocks Aluminum-rich minerals such as stone, corundum, andalusite, sillimanite, cordierite, etc.; in metaaluminous rocks, the feldspar-like, alkaline dark minerals and most aluminum-rich minerals in the above rocks do not appear, but feldspar appears. , quartz, common hornblende, common pyroxene, diopside and biotite, etc.
(3) The relationship between igneous rock formation conditions and mineral symbiotic combinations
The physical and chemical environment of igneous rock formation also has an important impact on mineral combinations. When magma cools in the deeper parts of the earth's crust, it is in an environment where the temperature slowly drops and the pressure is relatively high, giving it sufficient time to crystallize. Some of the minerals that begin to crystallize may be high-temperature types (such as feldspar, β-quartz), but as the temperature slowly drops, the high-temperature minerals formed early are no longer stable and gradually transform into stable minerals that adapt to low-temperature environments. For example, feldspar transforms into orthoclase, and β-quartz transforms into α-quartz. Plutons are therefore represented by the occurrence of low-temperature mineral assemblages. When magma erupts from the surface, the environment rapidly changes from high temperature and pressure underground to normal temperature and pressure. The rapid cooling of the magma has no time to crystallize and form a large amount of glassy or fine-grained high-temperature mineral combination rocks. At the same time, the high-temperature minerals that were previously crystallized underground when the magma erupted from the surface did not have time to transform into lower-temperature minerals, and still retained the structure of the high-temperature minerals. Therefore, the mineral assemblages of volcanic rocks are characterized by high-temperature minerals, fine-grained minerals, and vitreous minerals. In addition, in the high-temperature and high-pressure environment deep underground, primary minerals containing volatile components will be formed due to the participation of large amounts of volatile components in crystallization. The magma that erupts from the surface is difficult to crystallize hydrous minerals due to the large loss of volatile matter. Even the hydrous minerals such as hornblende and biotite crystallized deep underground by the magma are carried to the surface by the magma for oxidation. , dehydrate and decompose or partially decompose, and transform into other minerals such as magnetite and hematite, making the original minerals appear black or brown entirely or at the edges. This phenomenon is called darkening.
『撒』 The Yanshanian convective mantle was injected into the continent, the transformation of the continental crust of the eastern North China Craton and the formation of the Yanshanian orogen continental crust
(1) The Yanshanian convective mantle was injected into North China Eastern Continent
The development of sudden and intense igneous activity in the Jurassic-Cretaceous in the eastern part of the North China Platform shows that the Yanshanian convective mantle was injected into the continent, including the injection of a large amount of heat energy and material.
1. Huge heat fluxes (heatfluxes) are injected into the continent from the convective mantle
PublicAs we all know, the North China Platform was a typical craton before the sudden and intense igneous activity in the Yanshanian period. The geothermal calculations of the mineral inclusions in the diamonds of the upper mantle peridotite inclusions carried by the kimberlite magma of the Paleozoic North China Platform are equivalent to The heat flow value on the surface is 40~45mW/m2 (Zheng Jianping, 1999). In contrast to the eastern part of the North China Platform, the Ordos Loess Plateau has maintained craton structural stability during the Yanshanian and Himalayan periods, which can be used as a reference before the start of continental "revitalization" in the Yanshanian period in the eastern part of the platform. Based on the three heat flow values provided by the Shanghai Fengxian-Alxa Left Banner Geoscience Section (National Seismological Bureau Geoscience Section Editorial Committee, 1992) in the fixed edge area of the center of Ordos, the calculated average value is 44mW/m2, which is consistent with The heat flow values inferred during kimberlite eruptions are consistent. In this way, we can use the melting phase equilibrium diagram (Figure 2-64) of continental crust rocks (r-T1-υ-H2O) provided by Wyllie (1997) for discussion. The Yanshanian igneous rocks are mainly aluminous I-type granitoids, which need to reach temperature and pressure conditions that cause biotite to disappear (Bi-out) (Figure 2-64). The dehydration melting experiment of tonalite (T1) under 1GPa pressure (about 35km depth) (Rutter & Wyllie, 1998) shows that the solidus temperature is 825°C, and the Bi-out process occurs in the temperature interval of 850°C to 900°C. The amount of magma increases rapidly in the middle part, and the amount of magma can reach 22% at 900°C, leaving a refractory residue: Opx+Hb+Ga+Pl+Qz+Mt+Sp. The dehydration and melting experiment of biotite gneiss, which is widely distributed in Archean gneiss in eastern Hebei Province, under a pressure of 1GPa (Wu Zongxu et al., 1995) shows that the solidus temperature is 812°C, Bi-out is 837°C, and Hb- out is 887°C, and water-unsaturated granitic magma is formed in the interval of 812 to 950°C, with an amount of 20% to 30%. The remaining refractory residue is Opx+Cpx+Ga+Pl+Qz+Ru (rutile ). If we take the average thickness of the continental crust to be 35-40km and the shield temperature as the conditions for the initial uplift in the Yanshanian period, then, from Figure 6-7 we can see that under the conditions of 1GPa (the bottom of the crust about 35km thick), the initial The temperature of the continental crust is about 400°C. To produce ≥20% partial melt in the rocks at the bottom of the continental crust, the temperature must be raised to about 850°C, that is, the rocks at the bottom of the continental crust must be heated to about 450°C. (That is, initial 400℃ + temperature increase of 450℃ = 850℃).
Figure 2-64 Granite-tonalite-gabbro-H2O melting phase equilibrium diagram (according to Wyllie, 1997)
From convection The underplating of basaltic magma separated from the mantle into the bottom of the crust is the best mechanism for heating of the continental crust to induce local melting (e.g., Fyfe et al., 1973; Bergantz, 1989). BErgantz (1989) conducted a one-dimensional quantitative thermal simulation of basaltic magma underplating and local melting of continental crust rocks under 1GPa conditions. The temperature of the injected basaltic magma was 1250°C. The basaltic magma affected the overlying rock through its own cooling and crystallization. The heating of the earth's crust, the transmission of heat energy into the magma chamber and into the continental crust rocks are all considered to be conductive. The simulation results are that when the surrounding continental crust rock is tonalite (T1) and its initial temperature is 700 At ℃, the ratio of the total local magma produced to the total amount of basaltic magma crystallization is 0.4. For example, to produce 500km3 of magma, all 1250km3 of underplated basaltic magma needs to be cooled and solidified, and then the total amount of 500km3 of magma is required. Only 25% of the amount can be effectively separated from the melting zone and rise to form igneous rock. Therefore, only 125km3 of igneous rock can be formed. In other words, the ratio of the total amount of igneous rock formed to the underplated basalt magma is 0.1. From the above, if we only consider thickness (i.e. one-dimensional model), the formation of a 5km thick granite batholith requires the underplating of 50km thick basaltic magma and a continental crust melting zone containing a total local magma volume of 20km. supply. The melting experiment (Hirose & Kushiro, 1993) of lherzolite (KLB-1, and HK66) similar to the model upper mantle rock (pyrolite) (its chemical composition is shown in Table 2-11) under a pressure of 1GPa obtained 1250°C basaltic The melting degrees of magma are 6.5% and 17.9% respectively. In this way, if the melting degree is taken to be 10%, the generation of 50km thick basaltic magma needs to be supplied by a 500km thick convective mantle (if all local magma is to be separated). It should be noted that the above quantitative thermal simulation by Bergantz (1989) assumes that the initial temperature of the continental crust has reached as high as 700°C. From the beginning of the Yanshanian Period, the continental crust rock at a depth of about 35km was only 400°C. In other words, it must be The 400°C continental crust must first be raised by 300°C to reach the initial condition of 700°C. It is conceivable that heating the 400°C continental crust to 700°C requires the injection of a large amount of underplated basaltic magma into the continent. . In addition, many basaltic volcanic rocks and gabbro (υ) were formed during the Yanshan period. In other words, in addition to heating to melt the continental crust, the underplating basaltic magma was all consolidated (this part of the magma could not rise to form igneous rock bodies). In addition, a lot of basaltic magma is separated from the convective mantle and can rise to the shallows to form igneous rock bodies. Although we still need to conduct quantitative thermal simulation based on the properties of the Yanshanian igneous rocks in this area, the thermal simulation of Bergantz (1989) can be used as a guide, telling us that the Yanshanian igneous rock activity requires an extremely large amount of underplated basalt. The supply of magma, from a one-dimensional perspective, may require the supply of convective mantle with an interface with a depth of 670km. The basaltic Komati magma with an MgO content of up to 19% found in the checkerboard rock may be caused by this.It is isolated from the transitional mantle, which provides a support for the above quantitative results of thermal simulation. Obviously, the original cold craton lithosphere cannot provide such a large amount of heat flux into the continental crust. In other words, from the perspective of heat flux, the underplated basaltic magma will never flow from the cold lithosphere mantle. It must be separated from the underlying hot convective mantle (or asthenosphere).
2. The material (basaltic magma) separated from the convective mantle is injected into the continent
The convective mantle material injected into the continent refers to the basaltic magma separated from the convective mantle, which is often called primary ( or new) continental crustal materials (juvenile crustal materials). There are often two different views on the origin of Yanshanian basaltic magma, either from the asthenosphere (convective mantle) or from the lithospheric mantle.
From the perspective of petrological composition, the lithospheric mantle is the refractory residue after the convective mantle (or asthenosphere) separates basaltic magma. It is mainly harzburgite, and the asthenosphere is plump. The mantle peridotite (containing rich basaltic composition) is close to the original mantle composition. Ringwood (1975) proposed the Pyrolite model (upper mantle rock model) based on this (Table 2-10).
Table 2-10 Pyrolite model
(According to Ringwood, 1975)
In Table 2-10, oceanic tholeiitic basalt It is a pair of complementary products with harzburgite. 17% tholeiitic basalt and 83% harzburgite represent the composition of the parent body before the complements separated. Ringwood uses the prefix (pyr) of the English names of pyroxene and olivine. and ol) plus the suffix (ite) commonly used when forming rocks form an artificial rock called pyrolite to represent a model of the petrological composition of the original upper mantle, which is roughly equivalent to the composition of lherzolite in nature. Conversely, when 17% of the basaltic magma is separated from the pyrolite, 83% of the remaining basaltic magma is refractory harzburgite remnants. Compared with pyrolite, the residual harz. is rich in MgO, poor in Al2O3, CaO, and FeO. The weight percentage of FeO in Table 2-10 has almost no change, but the ratio of MgO to FeO changes greatly, indicating that FeO is relatively poor, while MgO is relatively Enrichment. Al2O3, CaO, and FeO are fusible components that enter the basaltic magma and separate away. MgO is a refractory component that remains in the peridotite. It has the same mineralogical characteristics as the disappearance of Cpx in the residual and the increase in the FeO component in the peridotite. adapt. It can be speculated that if harzburgite is to undergo melting again, the temperature at which the melting occurs must be higher, and the composition of the molten magma must be different from that of common basaltic magma (see Table 2-10).
Under high temperature and high pressure, the magma compositions produced by local melting of lherzolite and harzburgite are different (Hirose & Kushiro, 1993;Kushiro, 1990;Falloon & Danyushevsky, 2000), selected typical samples are listed in Table 2-11, and some oxide relationships are shown in Figure 2-62.
Table 2-11 Peridotite source rock and some selected magma compositions (all anhydrous 100%)
From Table 2-11 and Table 2-10 comparison, and Figure 2-65, it can be seen that the lherzolite and harzburgite source rocks used in the high temperature and high pressure experiment are similar to the pyrolite and harzburgite in the Ringwood model. Compared with the lherzolite In other words, harzburgite is rich in MgO and poor in Al2O3, FeO*, and CaO. In other words, it is the refractory residue after the original mantle peridotite separated the basaltic low-melting component. It can be seen from Table 2-10 and Figure 2-62: ① At the same high MgO content, the SiO2 of the magma melted by harzburgite is higher; under the same SiO2 content, the magma melted by harzburgite Magma MgO is higher (Fig. 2-65a); ② Under the same high MgO content conditions, the magma melted from harzburgite is poor in FeO. Under the same FeO content, the magma melted from harzburgite is rich in MgO. (Figure 2-65b); ③ Generally speaking, the Al2O3 of the magma melted by harzburgite is low. Under the same high MgO conditions, the Al2O3 of the magma melted by harzburgite is high (Figure 2-65c) ; ④ Under the same high MgO conditions, the magma melted from harzburgite has low CaO (Figure 2-65d).
Most magma in nature comes from lherzolite sources, and a small amount (such as Tonga, Troodos) comes from harzburgite sources (Table 2-12, Figure 2-66).
Most of the basalts and basaltic andesites with high MgO in the Yanshanian volcanic rocks in eastern North China are also evolved magmas. If separated olivine (Ol) is added, the MgO will be higher, and SiO2 will Reduced, they will enter the local magma range of lherzolite in the source area (Table 2-12, Figure 2-66); the chemical composition of some gabbro intrusions, such as chessboard rock and Shangzhuang intrusion , the MgO high is equivalent to the composition of komatiite-picrite-basalt, which in Figure 2-66 is equivalent to the local melting range of lherzolite as the source area (Table 2-12, Figure 2-66). In this way, the source area of the Yanshanian mantle-derived magma was not the mantle lithosphere dominated by harzburgite formed in the Archaean, but the plump convective mantle or asthenosphere dominated by lherzolite.
Figure 2-65 Comparison of melt products of harzburgite and lherzolite
Table 2-12 Selected natural mantle sources Magma composition (all 100% anhydrous, P2O5 not listed)
Continued table
Note: *Yu Jianhua et al., 1989, Plutogenesis and genetic evolution of gabbro in Beijing area.
Figure 2-66 harzburgiteComparison with lherzolite melt products
In this way, from the two aspects of the heat flux required for the formation of Yanshanian igneous rocks and the injection of mantle-derived magma, the Yanshanian convective mantle Injection into the continent is the driving force leading to the "revitalization" of the continent in the eastern part of the North China Craton.
(2) Crust-mantle interaction: mixing, differentiation, and delamination
1. Magma mixing: the main mechanism of mixing of crust-mantle materials
1)Zartman & Doe (1981) Lead (detection) structure (plumbotectonics) model (2nd edition). The lead (detection) tectonic model attempts to simulate the geochemical learning of U, Th, and Pb among major earth reservoirs (Figure 2-67, Figure 2-68). Figure 2-67 highlights the three long-term evolving reservoirs within the Earth, including the mantle, lower crust, and upper crust, and a short-term dynamic interaction of the above three reservoirs into the orogenic belt in an orogenic environment. interac-tion) to form the new crust of the orogenic belt (the column with black dots in Figure 2-67). Figure 2-68 shows the lead isotope evolution curves of the three long-term reservoirs predicted by the lead (detection) structural model and the new reservoirs in the orogenic belt formed by their short-term input. Their endpoints represent the modern isotope compositions of these four reservoirs. In this way, the formation of new crust in orogenic belts is the result of the mixing of material injected from the mantle with preexisting old continental crust material.
Figure 2-67 Structure diagram of lead isotope distribution reservoir in the mantle, upper crust, lower crust and orogen (according to Zartman & Doe, 1981)
Figure 2-68 A schematic illustration of the geochemical learning properties of U, Th, and Pb among major earth reservoirs that the lead tectonic model attempts to simulate (after Zartman & Doe, 1981)
2 ) Lead isotope composition of igneous rocks in the Yanshan orogenic belt. The lead isotope composition of the Yanshan orogen igneous rocks (SiO2 changes from about 45% basaltic to about 76% granitic) is significantly lower than the values of the four reservoirs in the Zartman & Doe lead (detection) structural model (1981) ( See Figure 2-67 ~ Figure 2-69) and Table 2-13, Figure 2-69.
Figure 2-69 Illustration of the lead tectonic model attempting to simulate the geochemical learning properties of U, Th, and Pb among major earth reservoirs (after Zartman & Doe, 1981)
< /p>Table 2-13 Comparison of Pb isotope composition
As can be seen from Table 2-13 and Figure 2-70, ① Mantle-derived checkerboard rock Komati The Pb isotope of lithic gabbro is obviously lower than that of the mantle and even lower than the global lower crust, indicating that it has been strongly mixed with the lower crust with low isotope values; ② The low value area of Pb isotope of Yanshanian igneous rocks is lower than that of northern Hebei -The lower crust of Liaobei Asia Province, indicating that the Pb isotopes in the source area of the lower crust are lower than those proposed by Zhang Ligang (1995). We useThe Pb isotope values calculated from the 9 feldspar lead in the metamorphic crystalline basement may be more reasonable to represent the source area of the lower crust. Its mixing with the mantle Pb isotopes can more reasonably explain the Pb isotope composition of igneous rocks; ③ The Pb isotope composition of Yanshanian igneous rocks records the magma Mixing and crust-mantle mixing; ④ Generally speaking, there seems to be a trend that the Pb isotope composition of the igneous rocks of J1 and J2 is greater than that of the igneous rocks of J3 and K1, which may indicate that the contribution of the mantle in the magmatic activity in the J1 and J2 periods is greater than that of J3 and K1. This seems to be consistent with the thermal simulation predicting that the crustal temperature at J3 and K1 is higher than that at J1 and J2. It is also consistent with the lithospheric surface delamination and surface magma activity during J3 and K1. Since the data are still sparse (a total of 23 (Parts), more data need to be accumulated, tested and revised; ⑤ For complexes formed in the same period, such as the Xuejia Shiliang Complex (except for the Heixiongshan granite), the isotope evolution range is narrower, generally As the SiO2 content increases, the isotope ratio decreases slightly (Figure 2-69), which is consistent with the model of magma mixing and crust-mantle material mixing. The obvious low value of Heixiongshan granite and the obvious high value of Sr isotope indicating that it comes from another source region.
Figure 2-70 SiO2-206Pb/204Pb diagram (A), SiO2-208Pb/204Pb diagram (B) of Mesozoic igneous rocks in Yanshan area
Figure 2-71 Decoupling of element and isotope exchange during magma mixing (according to Leisher, 1990)
3) Decoupling of element and isotope exchange during magma mixing. Through experimental studies of basic and acidic magmas, Leisher (1990) proposed the decoupling of element and isotope exchange during magma mixing. The exchange and mixing of isotopes is significantly faster than the exchange and mixing of elements (Figure 2-71). It can also be seen from Figure 2-71 that the exchange and mixing of incompatible elements is faster than that of the main elements.
The experimental results of Leisher (1990) are consistent with the magma mixing seen in nature. The microcrystalline diorite inclusions (MME) we see in granitic magma on field outcrops are direct evidence of magma mixing (Mo Xuanxue et al., 2002). It is generally believed that the magma mixing of granitic intrusions containing microcrystalline diorite inclusions is mainly physical mixing (or mechanical mixing) (magma mixing), and has not yet reached chemical mixing. It has a profound impact on the relationship between the inclusions and the host rock. It is correct in terms of the petrological composition and the chemical composition of the main elements, because the inclusions are often dioritic and the host rocks are often granitic. At the same time, microcrystalline (granular) inclusions are often considered to indicate its rapid Condensation solidifies, preventing chemical mixing from proceeding. However, the Sm-Nd isotope study of inclusions and host rocks (Pin, 1991) shows that the εNd (-3.2→-4.1) of the inclusions and the εNd (-3.5→-4.1) of the host granite are the same, which indicates that the isotopes Already sufficientexchange and achieve equilibrium, although there are still large differences in primary elemental and petrological composition. It can be seen from Figure 2-67 and Figure 2-68 that although the SiO2 content changes greatly, the change ranges of Pb, Sr, and Nd isotopes are small, especially the combinations formed by magma mixing in the same complex (such as Xuejia Shiliang The isotope composition of υ-δ-η-ξ) in the complex has been relatively fully exchanged and mixed. The Heixiongshan granite in the Xuejia Shiliang Complex does not belong to the above-mentioned mixing system, so its isotope composition is very different. (Figure 2-72). In this way, isotopes in igneous rocks can not only trace and identify different magma systems, but also better trace the mixing of crust-mantle materials and their relative contributions. If based on the principle of mass balance, according to Table 6-7, 206Pb/204Pb are 18.01 and 17.616 in the mantle, 14.576 in the crystalline basement of the continental crust, and 16.369 in the Yanshanian igneous rock, then the respective contributions of the mantle and lower crust during mixing are 52.2% and 47.8%.
Figure 2-72 Illustration of the lead isotope composition of the Xuejia Shiliang ring rock mass
If 143Nd/144Nd is used to estimate, take the Yanshanian Jiang The two average values of the Miaohe Yanghutang gabbro are 0.5126385 (εNd=+1.35), the crystalline basement value of the North China continental crust is 0.511796 (Guo Jinghui, 2000), and the Yanshanian igneous rock is 0.511927. The estimated contributions of the mantle and lower crust are 15.5. % and 84.6%.
4) Yanshanian igneous rock structural combination. The Yanshanian igneous rock combination is the volcanic rock combination of basalt-potassic trachybasalt-potashite-andesite-trachyte and trachydacite-rhyolite, and the corresponding gabbro-monzonitic gabbro Intrusive rock combinations of rock-monzonite-monzonite-syenite and quartz monzonite-granite (see Section 3 of this chapter). Taking the volcanic rock combination as an example, as shown in Figure 2-73, it is classified according to SiO2 content, forming a normal distribution (or unimodal), with andesitic (A) as the peak. For example, according to TAS classification, the andesitic content in the study area Quality (A) should be trachyandesite; according to SiO2-K2O classification, the volcanic rocks in the study area are mainly HKCA series, so they are typical orogenic belt igneous rock combinations; compared with the Andes, they lack the MKCA series and its important members of andesite (amphibite). andesite (monzonite) and trachyte (syenite) and trachydacite (quartz monzonite) of the HKCA series (Deng Jinfu, Liu Houxiang et al., 1996).
The above-mentioned igneous rock structural combination shows that the mixing of crust and mantle materials formed the HKCA series with andesite as the peak. If the Archean continental crust is composed of T1T2G1G2, then in the Yanshan period due to the convective mantle materials ( The injection of basaltic magma) "basicized" the continental crust at that time. Therefore, it moved in the direction of andesite with low SiO2 content, and at the same time it became relatively high in K2O (high alkali). ThisThis is because the Archean continental crust, which is mainly MKCA series and is low K2O (low alkali) T1T2G1G2, experienced local melting and separated a large amount of high K2O (high alkali) granitic-syenite magma. Judging from the TAS diagram and the SiO2-K2O (Figure 2-46) relationship, the primary magma or nearly primary magma separated from the convective mantle belongs to the MKCA series, and the K2O and alkali are not high. Therefore, the entire igneous rock series The main body enters the HKCA series, and the increase in K2O and alkali is mainly related to the melting of the continental crust.
2. The main mechanism of crust-mantle material differentiation: local melting and AFC model
1) Partial melting: Whether it is the original Archean T1T2G1G2 and mafic Mafic metamorphic rocks can undergo local melting due to later heating, or after the felsic magma formed by underplating basaltic magma in the Yanshanian period or mixed with continental crust and solidified into rocks, local melting can occur. This process forms magma. At the same time, complementary refractory residues must be left behind, thus leading to the differentiation and evolution of crust-mantle materials. If the local melting degree is similar to that of the source rock, but the pressure is different, the mineral combination of the resulting magma and its residue will be different.
Figure 2-73 Comparison between the Tibetan Plateau continental collision volcanic rocks, East Yanshanian volcanic rocks and Andean arc volcanic rocks
Figure 2-74 is composed of basaltic Amphibolite melting phase diagram and residual mineral assemblage (Wyllie et al., 1997); Figure 2-75 shows the liquidus-equilibrium mineral assemblage of trachytic (syenite) magma formed by local melting of felsic rocks , the latter is the residual mineral assemblage (Deng et al., 1998); Figure 2-76 shows the mineral assemblage in equilibrium with the liquidus line during the melting experiment of Orzogranite (T2), the latter also represents the residual mineral assemblage (Laan & Wyllie, 1992).
A common feature of Figures 2-74 to 2-76 is that the residual mineral combination left behind when local melting of continental crust rocks occurs at lower pressures (<1.4~1.7GPa). The granulite phase combination containing plagioclase is an eclogite phase combination without plagioclase under high pressure (≥1.4~1.7GPa, about ≥55km thick continental crust). Hb is often present when the temperature is relatively low. At high times, there are only garnets.
2) Assimilation-fractionational crystallization (AFC) model: When the underplated high-temperature basaltic magma heats the continental crust, assimilates continental crust rocks, or mixes with magma caused by local melting of the continental crust, , the basaltic magma itself must undergo strong crystallization. Chemical characteristics of Yanshanian igneous rocks, including AFM diagram (Fig. 2-47), SiO2-K2O diagram (Fig. 2-46), TAS diagram (Fig. 2-45) and SiO2-(Na2O+K2O-CaO) diagrams (Figure 2-48) all show that the AFC model is one of the main mechanisms of magma differentiation and evolution, that is, the primary basaltic magma separated from the convective mantle is rich in MgO, low in potassium, subalkaline, and CA (according to Peacock's According to the classification of alkali-calcium index), the felsic magma of continental crust origin is relatively FeO-rich, high-potassium, alkali-rich, and AC (according to the classification of Peacock's alkali-calcium index), forming an overall structure through the AFC mechanism. The igneous rock series with little change in the FeO/MgO ratio, high potassium (HKCA), relatively rich in alkali and AC (according to Peacock's alkali-calcium index classification) can also be inferred from Figure 2-74 to Figure 2-76. Crystallization The separated mineral combination is granulite phase combination under low pressure and eclogite phase combination under high pressure.
Figure 2-74 Dehydration melting phase diagram of amphibolite (according to Wyllie et al., 1997)
Figure 2-75 H2O unsaturated Andesite liquidus surface (according to P.J. Wyllie, 1976, quoted from Deng Jinfu, 1987)
Figure 2-76 H2O saturated Nuke gneiss liquidus surface (according to P.J.Wyllie, 1992)
From the above, the interaction between crust and mantle materials is the mixing of materials on the one hand and the differentiation and separation of materials on the other hand through magma.
3. Basaltic eclogite-induced lithospheric delamination
The National Plan for Research on Continental Dynamics in the United States (Xiao Qinghui et al., 1993) pointed out that new lithosphere delaminations formed in island arcs The overall composition of the crustal material is equivalent to basalt, but the velocity structure of the continent shows that the average composition of the existing continental crust is closer to andesite. Perhaps a slightly more siliceous-aluminous crust has formed in the island arcs in the past, or perhaps some process has long removed the denser components from the continental crust. The delamination of part of the continent's lower crust and lithospheric mantle during the collision has been suggested as an explanation worthy of investigation. If this effect occurs regularly, it could have profound implications for understanding the structure and composition of continents.
The above discussion tells us that ① the injection of convective mantle material into the continent adds a large amount of material to the continental crust. Due to the addition of a large amount of basaltic magma, the original Archean felsic continental crust must be Evolve towards a more mafic direction; ② At the same time as the crust and mantle materials are mixed, the crust and mantle materials differentiate and separate through magma (such as local melting, AFC process, etc.), and generally differentiate into more felsic materials. Magma, separates leaving remnants of a more mafic mineral assemblage. However, current geophysical detection shows that the average vP of the Ordos block continental crust is 6.3km/s, and that of the Taihang and Yanshan orogenic belts is 6.2~6.3km/s (Jia Shixu, 2003, unpublished, internal report). The former is roughly equivalent to The composition of granodiorite (G1) (average SiO2 is 65%), which is roughly equivalent to the composition of granodiorite (G2) (average SiO2O2 is 70%). The Ordos block has maintained a structurally stable platform (craton) since the formation of the Archean crust. Therefore, the composition of the continental crust can be used as a reference for the composition of the original continental crust before the Yanshanian "continental activation" in the eastern part of the North China Platform. . There is basically no distribution of Cenozoic basalt in the Yanshan orogenic belt. It can be generally regarded that there is no addition of mantle material in the Cenozoic. In addition to the possible thinning of the crust, its continental crust can generally be regarded as a product of the Yanshan movement. Therefore, it can be seen that the Yanshan Movement "acidified" the continental crust, which is in obvious contradiction with the "basic" continental crust made by the addition of a large amount of basaltic magma. It must require a considerable part of the mafic to return to the mantle. , the demolition effect may be the best explanation.
The delamination of the lower crust plus mantle lithosphere that “acidifies” the continental crust must meet the following conditions: ① The density of the detached lower crust (+mantle lithosphere) must be greater than the underlying soft crust For the Yanshan orogenic belt, there is a special problem that the original Archean lithospheric mantle is low-density (mainly harzburgite); ② The main body of the detached lower crust should be magnesium Ferrous and ultramafic. The best object that satisfies both conditions is the occurrence of large amounts of eclogite.
As discussed earlier, during the evolution of crust-mantle interaction, felsic magma was continuously formed and mafic and ultramafic mineral combinations were separated. This part of the mafic and ultramafic residual mineral components in the early evolution of the orogenic belt (J1 and J2) may transform into the eclogite phase under the thickening of the continental crust driven by secondary contraction structures in the late J1 and late J2 periods. The rocks, together with the early underplating basaltic magma, may have mostly cooled and solidified into rocks during the heating process of the cold continental crust. They will also transform into the eclogite phase under the conditions of thickening of the continental crust. The presence of no or weak negative Eu anomalies and high Sr/Y ratios as well as andesite (monzonite)-trachyte (syenite, quartz monzonite) with high K2O in the J3 and K11 igneous rock assemblages indicates that at that time, It is a thickened continental crust (≥55km), and the residual mineral combinations consolidated and separated from the underplated basaltic magma should be eclogite phase rocks. The linear distribution of igneous rocks at J1 and J2 and the planar distribution of igneous rocks at J3 and K11 indicate that the volume of eclogite phase rocks accumulated in the thickened lower crust and lithospheric mantle at the beginning of J3 is large enough (or has reached a critical value ), resulting in large-scale deconstruction of the lithosphere.
(3) The formation of the continental crust of the Yanshanian orogenic belt
Except for a small amount of Cenozoic sedimentary rocks, most of the rocks exposed on the surface today are igneous rocks formed in the Yanshanian period. (including volcanic rocks and intrusive rocks), metamorphic rocks and sedimentary rocks, as well as modified pre-Jurassic rocks involved in the Yanshan movement. Many large granitic batholiths, such as the Badaling and Dahenan batholiths, are representatives of the middle and lower parts of the upper crust exposed on the surface. Yunmengshan granitic gneiss rock base, Shatuozi eye-shaped gneiss body, Wudaohe granitic eye-shaped gneiss body, Shicheng dioritic gneiss body, Changyuan dioritic gneiss body The body and the source rock are the Great WallThe Sihetang amphibolite-phase metamorphic rocks of the Jixian Series are the orogenic root zone rocks exposed on the surface and represent the middle crustal rocks of the orogenic belt (Davis et al., 2001; Beijing Geology, 1991). Davis et al. (2001) pointed out that it is strange that zircons from dioritic gneiss intrusions do not contain inherited components, but, unlike young diorite rocks, granitic gneiss and eyeball Zircons from gneiss-like rocks contain inherited components, and their upper intersection points range from 1.7 to 2.4 Ga. The SHRIMP study of this project supports the conclusion of Davis et al. that the dioritic gneiss intrusion has no inherited zircons, while the Yunmengshan granitic gneiss intrusion has inherited zircons (2.4Ga). In addition, granulite phase plagioclase pyroxenite inclusions (120-140Ma) and eclogite phase garnet pyroxenite (with pile crystal structure) found in the Cenozoic Hannoby basalt (Fan Qicheng et al., 1998, 2001), indicating that there were indeed newly formed granulite-phase rocks in the lower crust and eclogite-phase rocks formed in the mountain root zone during the J-K period, supporting the theory that there was a thickened continental crust and basaltic magma underplating in the Yanshanian period. model (Deng Jinfu et al., 1996). The magma rose through the Archean and Paleoproterozoic caprock. Showing that they are clearly new organisms that branched off from the Earth's mantle.
The formation process of the continental crust of the Yanshanian orogenic belt can be roughly summarized as follows:
1) The injection of heat and material from the convective mantle was the result of the growth of the continental crust and the strong modification of the ancient continental crust. The driving force and injection mechanism are mainly the underplating of mantle-derived basaltic magma (basaltic magma underplating).
2) The formation and transformation of the continental crust are mainly achieved through igneous processes. The formation and evolution, migration and positioning of magma are the basic processes of the growth of the continental crust.
3) Crust-mantle interaction is the key mechanism for the formation and evolution of continents. The magmatic process is the main mechanism for implementing crust-mantle interaction, which includes the mixing of crust-mantle magma, the differentiation of magma and its interaction with Separation processes of residual mafic mineral assemblages and delamination of mafic and ultramafic eclogite phase rocks.
4) The formation of thickened continental crust driven by the shrinkage structure of the orogenic belt is necessary for the formation of a large number of eclogite phase rocks, and is necessary for the detachment of the lower crust and lithospheric mantle.
5) The formation and transformation of the continental crust not only involves the addition of convective mantle materials, but also the return of crustal and lithospheric materials to the convective mantle. These two mechanisms are responsible for the maturation of the continental crust, and ultimately Necessary for the formation of mature continental crust dominated by granitic matter.
6) The sequence of magmatic tectonic events and the combination of igneous rock structures are key records for tracking the process of convective mantle injection into continents until the formation of mature continental crust.
7) The joint constraint of Pb isotope and Nd-Sr isotope composition is the best tracer record of the growth and transformation of continental crust. It is not only possible to identify completely new continental crust (represented by Archean TTG), Or is it a modified continental crust that is re-injected by the convective mantle (Taking the Yanshanian G as a representative), it is also possible to invert the relative contribution of the mantle in the modified continental crust and the original ancient continental crust. Therefore, the emphasis on isotope composition is because, ① the petrology and main elements of igneous rocks, due to the delamination of a considerable part of the basaltic eclogite back to the mantle, the composition of the continental crust is generally only equivalent to the long-term composition of the continental crust left after the injection and modification of the convective mantle. ② In the crust-mantle interaction, the exchange rate of isotopes is much faster than that of main elements and trace elements, which may represent the relative contribution in the mixing of crust-mantle materials; ③ Compared with the main elements and For trace elements, the reference values of isotope ratios of end-member members of crust-mantle materials are easy to determine, so calculating their relative contributions is more reasonable and operable.
『8』 Large-scale magmatic event and melting mineralization in the Proterozoic Qilian ancient continent block
(1) Large-scale magmatic event in the Proterozoic Qilian ancient continent< /p>
The Jinchuan deposit is famous for its huge Cu-Ni-PGM reserves. As an exposed area of only 1.34km2, the nickel metal reserves reach 546×104t, and the rock mass mineralization rate is as high as 60%. Independent ultramafic rock bodies are unique in the world. In the face of such highly concentrated mineralization facts, the deep melting of the deposit, or the mechanism of immiscibility between large-scale sulfide liquid phase and silicate melt before invasion, is definitely present, and the final ore-forming magma The room is actually a slurry room, where the melt extremely rich in sulfide liquid phase condenses and crystallizes in place to form rocks and minerals. This understanding has been recognized by Chinese scholars for a long time and has been continuously enriched (Tang Zhong et al., 1995). The concept of "small rock bodies forming large mineralization" proposed by Academician Tang Zhongli is a profound revelation of this mineralization mechanism. Such a large-scale metal sulfide accumulation must be possible only if a larger-scale magma source is provided. This is a common concern in the study of magmatic Cu-Ni-PGE sulfide deposits in the late 1990s (Keays, 1997; Pirajno, 2002). Whether a large amount of sulfur comes from the mantle or the crust, the results of sulfur isotope research show that the answers are diverse. There are mantle sources, such as Jinchuan et al. (Tang Zhong et al., 1995; Li Wenyuan, 1996), and crustal sources or added crustal sources, such as Noril'sk et al. (Naldrett et al., 1996). It has been deeply recognized that sulfur saturation is the main reason for the immiscibility or dissolution of sulfide liquid phase (Naldrett, 1989; Rad'Ko, 1991; Brugmann et al., 1993; Keays, 1995), and reducing sulfur The factors of saturation in magma, the addition of crustal materials, magma mixing, etc. are the main factors (Ripley and Li, 2002; Ripley, 1981; Naldrett et al., 1993, 1989, 1999; Lambert et al., 1991, 1999 , 2000)). But it is certain that metallic nickel comes from magma, and a large amount of sulfur alsoMainly derived from magma, and magma is mainly derived from mantle sources. For magma to provide super-scale nickel accumulation, it is possible only if there must be super-scale magma. In the continental crust environment, the most likely form of the existence of supernormal-scale magma is large igneous provinces (LIPs) related to mantle plumes (Lightfoot et al., 1997; Keays, 1997). Because it only originates from the "D" layer below the core-mantle boundary (CMB) (in seismic terms, it is the thermal and chemical interaction zone between the core and the mantle, the transmission zone of thermal energy originating from the core), and the mantle plume is believed to originate from this zone. Only the mantle plume of Lay et al., 1998) may cause sudden supernormal thermal events in local areas and produce large-scale magmatism to form LIPs, and provide a large amount of mineralization components S, Cu, Ni and PGE.
The most important feature of the existence of LIPs in continental environments is continental overflow basalt (CFB). Currently, LIPs with large areas of CFB distribution have been confirmed to include the Siberian Platform (Noril'sk deposit) in Russia and the Central Continent in the United States. The Mid-continent Rift System (Duluth, Me11en deposit), the Karoo igneous province in South Africa, the Deccan Plateau in India and the Emei basalt in southern China, etc., Pirajno (2002) among others The main characteristics of large igneous provinces related to mantle plumes are summarized in the book "Ore Deposits and Mantle Plumes" as follows (Figure 4-45): The mantle plume originating from the "D" layer rises to the bottom of the lithosphere. Due to the decompression mushroom Deep melting occurs at the head of the mushroom-shaped mantle plume, and lithospheric fragments are sucked in. As it passes upward through cracks in the crust, the magma filters to form a high-level magma reservoir in the crust; some of the reservoirs reach the surface and erupt to form continental flow basalt (CFB). or volcanic rocks, and the rest consolidated in situ (magma chamber) to form layered igneous complexes. The high Mg of melts initially derived directly from the mantle plume axis contributes to the early phase of overflow basalt, and the migration of late mantle plume melts may erode the thermal boundary layer of the lithosphere and carry lithospheric fragments back to the mantle. Therefore, mantle plume melts are enriched in incompatible elements compared with oceanic crustal lithosphere. In short, LIPs formed by mantle plume-lithosphere melting mainly exhibit three-dimensional characteristics: the first is CFB; the second is layered igneous intrusions at high levels in the crust; and the third is basic dyke groups. (mafic dyke swarms). The three are interrelated and cause and effect each other. In particular, CFB is closely related to large-scale basic and ultrabasic layered intrusions. The latter is a magma chamber or a channel that supplies overlying basalt. Therefore, in terms of macroscopic and direct manifestations, CFB is the most typical manifestation of LIPs.
Figure 4-45 Mantle plume and rockSchematic diagram of the interaction between the lithospheric mantle, the mafic-ultramafic magma reservoir at the bottom of the crust, the emplaced continental flow basalt (CFB) and the corresponding bedrock complex
Due to the exposure of CFB to The surface of the Earth is the most thoroughly studied. Statistical results of 14 CFBs around the world, all developed from the Archean (2772Ma) to modern times (15Ma), with a development period of 1 to <10Ma, typically tending to be Fe-rich, mainly composed of continental tholeiitic basalt, and the bottom of the Archean CFB There are komati rocks produced, and in the Proterozoic CFB, there is a tholeiitic basalt province in the Bangemall Basin in Western Australia that was formed at 1.6 Ga, which is very close to the age of Jinchuan. In terms of time and space, CFB is closely related to the uplift and expansion of the earth's crust, so it can be formed in the midcontinental rift system (MCR in the United States, located in Lake Superior and Kansas, the famous Duluth layered complex, namely Located in it, the origin is the same as the basic lava distributed over a large area), and can also be produced in volcanic rift valleys on the continental margin (North Atlantic Igneous Province). The basic dyke group is related to the breakup of the continent. Its radial basic dyke group may reflect the central position of the mantle plume. A single dyke is from a few meters to 200m wide and hundreds of meters to 1000km long. The basic dyke group is the continent Reconstructed landmark remains. Layered igneous rock intrusions are actually a general term. A large number of different types of intrusions in large igneous provinces can be included in its category. They are the main carriers of magmatic Cu-Ni-PGE deposits.
Based on the huge accumulation of metal content in the Ni-Cu-PGE-bearing ultramafic intrusion in Jinchuan, it is inferred that it is the result of the mantle plume and is a hot spot in geological history. It is hypothesized that the Qilian Mountains were ancient in the early Mesoproterozoic LIPs formed by the mantle plume on the continent must have layered igneous rock intrusions (the Jinchuan intrusion can be regarded as one of the ore-forming rock bodies) and corresponding distribution of CFB and basic dyke groups.
(1) The extremely thick volcanic rock series distributed in the Zhulongguan Group of the Jingtieshan micro-block in the western section of the North Qilian (see Chapter 3 for details) can be considered to be a local exposure of CFB or Residue. Xia Linqi et al. (1999, 2000) have demonstrated in detail the characteristics of CFB from the perspective of volcanic rock petrology and geochemistry, and will not go into details here. It is believed that CFB must be distributed over a wide area. The limited area found today can be explained by the fact that the Qilian ancient continental block where it was formed after the formation was dismembered and reassembled and uplifted, and suffered denudation to varying degrees. The Changcheng Period metamorphic basic volcanic rock series discovered in the western section of Longshou Mountain should also be the exposure of CFB during this period. Due to the lack of high-quality age data and detailed volcanic rock petrology research, it is only speculation at present, but Zangbutai, The Qingshiyao ultramafic rock is likely to be komatitic or picrite lava, which can form a complete CFB combination with the continental tholeiitic basalt and alkaline basalt in the Zhulongguan Group.
(2) The basic dyke group has not yet been reported in the Qilian Mountains. However, a large number of gabbro dykes found in the CFB distribution area in the Zhulongguan Group of the Jingtieshan microblock can be Considered as a local exposure of the basic rock wall group, BeiqiA large number of irregular basic rock dykes were also found in the 1:50,000 regional geological survey of the western section. In addition, a large number of gabbro and diabase veins distributed parallelly in the Mesoproterozoic dolomite in Longshou Mountain are also basic dykes worthy of re-understanding their geological significance. Since the spatiotemporal distribution and geological significance of Precambrian mafic dykes in the Qilian Mountains have not been examined on a larger scale in the past, the existence and characteristics of mafic dyke groups are issues that need urgent investigation and research.
(3) The layered igneous rock intrusions in the Qilian Mountains have been the object of most attention in previous studies. However, the mineralization investigation of the mafic-ultramafic intrusive bodies in the Qilian Mountains (including Longshou Mountain) has only been conducted so far. It was discovered that outside the large-scale mineralization rocks in Jinchuan, no mineral deposits were found in more than 10 mafic-ultramafic intrusions distributed in Longshou Mountain. However, in the Paleoproterozoic basement Hualong Group in South Qilian, Lala was found. The most noteworthy thing about small magmatic Cu-Ni-PGE deposits such as Shuixia (see Chapter 4 for details) is that although the metal reserves are not large, they are whole-rock mineralized, and the rock mass is the ore body, which is of unique significance. If we understand the existence of the Jinchuan deposit from the perspective of the early Mesoproterozoic LIPs of the Qilian Mountains ancient continental block, the Lashuixia ore-bearing rock mass and the Jinchuan rock mass are connected, and they belong to the same mantle plume that partially melted and intruded the bottom of the lithosphere. The result of penetration mineralization. It is predicted that there may also be hidden layered intrusive bodies, which are the geological prerequisites for discovering new mineralized rock bodies.
The author proposed the idea of large-scale mantle plume large igneous provinces (LIPs) in the early Mesoproterozoic in the Qilian Mountains, aiming to explore the material source and geological implications of this world-class mineral deposit in Jinchuan. And try to explain the accumulation of huge amounts of metal materials in the upper crust of the earth's mantle. Due to the complexity of the content contained in this topic, only a framework conjecture is proposed, but there are some unavoidable issues. The first is the age issue, the current controversy over the age of the Jinchuan rock mass (Mesoproterozoic/Neoproterozoic), the isotope test age of the CFB large span in the Jingtieshan microblock, the diagenetic and mineralization age of the Lashuixia rock mass, etc. All of them restrict the clear grasp of the entire magmatic-tectonic event; secondly, because the scope of magma is the geological range of the late orogenic belt and its marginal blocks, the rapid and multiple structural changes in the orogenic belt have transformed the early igneous rock events beyond recognition, and supporting Recovery is extremely difficult. But in any case, reshaping the Precambrian tectonic-magmatic events in the Qilian Mountains is of great significance to the systematic understanding of the mineralization evolution process of the Jinchuan deposit and further development of regional prospecting deployment.
(2) Sulfide deep fusion mineralization of the Jinchuan super-large world-class deposit
As mentioned before, the degree of sulfur dissolution in the silicate melt mainly depends on The content of FeO, followed by the content of CaO, MgO and Na2O. According to Godlevsky's (1981) silicate melting experiments on Norsk ore-bearing rocks, sulfide exists in a suspended state at high temperatures (1400-1500°C) and the content in the silicate melt can reach more than 15%. Ma-cLean (1968) found that there are a large number of immiscible phases in the bulk FeO-Fe3O4-SiO2-FeS system (Figure 2-14). He pointed out that during the crystallization process, a homogeneous silicate melt containing a small amount of sulfur may saturate the sulfur to form sulfide, resulting in the existence of a molten sulfide liquid phase. The crystallization trajectory depends on the original chemical composition and the decrease, equilibrium, or increase in oxygen fugacity during crystallization. In the FeS-FeO-Fe3O4-SiO2 system, the lowest temperature at which the two liquid phases exist is 1140°C. The solubility of sulfur in this simple basic magma may be about 4%, and the silica in the sulfide/oxide liquid phase is about 1%. The actual solubility of sulfur in nature can be as high as 15% (Godlevsky, 1981). Crystallization of this immiscible sulfurized liquid phase begins only after complete crystallization of the silicate. The genetic mineralogy study of the Jinchuan intrusion shows that the Ni content in chromite is depleted (Barnes et al., 1999), sulfide daughter minerals are rare in olivine magma inclusions (Yang Xuanzhu et al., 1991), and immiscibility affects silicon. before crystallization of acid salt minerals.
Natural basalt melt contains only 0.03% sulfur by weight (MacLean, 1968), which is 100 times lower than the maximum value in the FeO-Fe3O4-SiO2-FeS system, while the natural magma contains even lower FeO and silicon A large amount of CaO, MgO and Al2O3 in the acid salt combine with FeO, greatly reducing the concentration of FeO, resulting in a reduction in the solubility of sulfur in the melt. Once the sulfide melt separates from the silicate melt, the sulfide-rich melt can easily form massive ores. In the early stages of crystallization, Ni and Cu replace Fe into the silicate structure (clinopyroxene and olivine) because the N-O bond is stronger than the Fe-O bond. Before the sulfide melt melts away, the higher the degree of silicate crystallization and the higher the oxygen fugacity, the more Ni tends to enter the silicate mineral and rarely combines with the sulfide melt. Therefore, due to the high mineral content of the Jinchuan deposit, there will be no obvious crystallization of olivine before the sulfide is melted away, otherwise it will be difficult to form a rich ore.
Naldrett (1989) believed that the differences in the material composition of different mineral deposits are due to the separation and branching equilibrium of sulfides. The sulfide dissolved in the magma becomes an immiscible liquid and melts away due to the assimilation of silicon-rich materials by the mafic-ultramafic magma. Different proportions of sulfide separation are the result of varying degrees of assimilation. The sudden melting and sinking of the sulfide liquid phase may be due to the mixing of silicon-rich magma. It can be seen that magma mixing is important to the mineralization of magmatic sulfide deposits. Lambert et al. (1988, 1989) used rare earth elements, Sm-Nd isotopes and Re-Os isotope geochemical studies, combined with field and petrological characteristics to determine that there are at least two end-member magmas in the Stillwater complex, one is U-shaped Magma is a basalt partially melted by the Wyoming Archaean quasi-continental lithospheric mantle; the other type is A-type magma, which is a tholeiitic basalt formed by the partial melting of mixed basaltic magma in the upper crust and mafic-ultramafic rocks in the lower crust. . Think Stillwater ComplexThe relatively wide γOs and εNd values in the product area are the result of the geochemical mixing of different magmas, which plays a role in the entire history of the rock-forming and mineralizing magma chamber. Other foreign magmatic sulfide deposits also have this characteristic, such as Noril'sk picrite and PGE-rich Ni-Cu sulfide ore, γOs=+4.1~+14.2 (Lambert et al., 1989), Bushveld complex Rock UG-2 chromitite and Mequimes layer, γOs=+33～+68 (Hart and Kinloch, 1989), Sudbury igneous complex, γOs=+322～+854 (Walker et al., 1991) . These data imply that some degree of crustal material was added to the system at each stage. The characteristics of εNd<0 of the Jinchuan rock mass (Li Wenyuan et al., 2004) and γOs=+9.1~+122.6 (Liu Minwu, 2004) also reflect the characteristics of the addition of crustal materials. It is the sulfide liquid phase caused by sulfur saturation before crystallization of the Jinchuan rock mass. The main reason for large-scale melting or immiscibility. Figure 4-46 can illustrate its mineralization process.
When the ancient mantle plume reaches the bottom of the lithosphere, it undergoes partial melting due to decompression and produces a large amount of magma. Lithosphere fragments are continuously added to the magma, so the composition of the rising magma (1) will change; Rising to the boundary between the lithospheric mantle and the crust, a large-scale magma chamber (or magma reservoir) is formed. At this time, the magma is still unsaturated with sulfur. As the crustal material in the magma chamber is brought in, contamination and new magma are The injection causes magma mixing (2), saturating the sulfur in the magma chamber, and causing large-scale immiscibility between the sulfide liquid phase and the silicate magma melt (deep melting of the sulfide liquid phase) to occur (3), Sulfur-loving metal elements such as Ni, Co, Cu, and PGE enter the sulfide liquid phase one after another, causing the basic silicate magma melt in the upper part of the magma chamber to be depleted of S, Ni, Co, Cu, and PGE; when the magma is depleted of metal elements, While the upper intrusion pierces (melts) the earth's crust and erupts and overflows on the surface to form basic lava (5), a large number of basic rock wall groups will also be preserved as magma conduits. More importantly, the central and lower parts of the deep magma chambers will be preserved. Sulfide-containing magma or directly sulfide liquid phase intrudes and penetrates upward due to the extrusion of the earth's crust, forming a high-level sulfide-containing magma chamber (or pulp chamber) in the earth's crust (4). Such sulfide-containing magma There should be multiple chambers, but the size and the amount of sulfide contained may be different, and each magma chamber may be injected with magma and ore more than once. Eventually, magma crystallization accumulated in the high-level magma chamber and the sulfide liquid phase was located at the bottom of the magma chamber, forming a layered intrusion. Later tectonic changes caused the spatial location of the ore-bearing layered intrusions to change, standing steeply or lying sideways and denuded to expose the surface or be located in the shallow part of the crust, becoming available ore deposits or potential deposits.
Figure 4-46 Schematic diagram of the sulfide melting and mineralization evolution of the Jinchuan magmatic Cu-Ni-PGE deposit
The Jinchuan deposit is the world's largest magmatic Cu-Ni-PGE deposit. Its mineralization background and deposit origin have always been issues of interest to people. Foreign researchers of magmatic Cu-Ni-PGE deposits have also shown great interest. But in general, the introduction to the outside world is not enough. In the 100 years since the discovery of the Sudbury deposit, new understandings have emerged one after another. The famous "Economic Geology" magazine has published four special issues in 1971, 1990, 2000 and 2002. The academic issues raised by research are becoming more and more widespread. Since the Jinchuan deposit plays an important role in China's magmatic Cu-Ni-PGE sulfide deposits, the understanding of its mineralization geological background and the origin of the deposit involves the research of similar deposits throughout China. The views on magma mixing and mineralization, the addition of crustal materials, etc. are all issues of great concern to everyone. The petrochemistry of the Jinchuan rock mass has the chemical composition characteristics of komatiite, and the ore geochemistry is characteristic of tholeiitic basalt deposits. It is speculated that one of the end-member magmas is the komatiite magma that has been contaminated by the crust (based on the magma in the eastern section of the rock mass (domestic chamber), the (Pt+Pd)/(Os+Ir+Ru) ratio is low, the PGE content itself is low, and only Ni-Cu ores are formed; the magma chamber in the western section is tholeiitic basalt magma to form Ni-Cu Characterized by sulfide ores and Pt and Pd enrichments. This understanding is only a speculation and requires more detailed research.
In addition, the contribution of the Jinchuan deposit's composite hydrothermal processes to mineralization is also significant, and there may be multiple stages of activity, especially in the successive tectonic transformations in the later stages of diagenesis and mineralization. Various high-temperature composite hydrothermal processes The activity of liquid is very important for the formation of copper-rich massive ores and the local enrichment of PGE, and it is necessary to conduct in-depth research.
『九』 Mineral chemical characteristics and metamorphic temperature and pressure conditions
The main minerals of granulite xenoliths are pyroxene, plagioclase and biotite, and their electron probe The data are listed in Table 4-17. The data were mainly analyzed by the Electronic Probe Laboratory of the Key Laboratory of Orogens and Crustal Evolution of Peking University, Ministry of Education. The analytical instrument was EPM-810Q. Test conditions: accelerating voltage 15kV, beam current 3.8nA. A small amount of data was analyzed by the electronic probe laboratory of the Institute of Geology and Geophysics, Chinese Academy of Sciences. The analytical instruments and test conditions are the same as in the previous section.
Table 4-17 Average chemical composition (wB/%) and metamorphic temperature of the main rock-forming minerals in basic granulite xenoliths
< p>Continued table
Note: FeO* is all iron; (a) was tested by the electronic probe laboratory of the Institute of Geology, Chinese Academy of Sciences, and the rest were analyzed by the electronic probe laboratory of the Department of Geology, Peking University ; TWB is the result calculated using the Wood-Banno (1973) lherzolite thermometer, and TW is the result calculated using the Wells (1977) lpyroxene thermometer.
1. Pyroxene
Pyroxene is one of the most important rock-forming minerals in granulite xenoliths. According toAccording to the classification scheme of Deer (1978), clinopyroxene is classified as pyroxene, while clinopyroxene belongs to ordinary pyroxene and diopside (Fig. 4-24).
Figure 4-24 Wo-Di-Hd-Fs-En composition diagram of pyroxene in granulite xenoliths
(According to Deer, 1978)
The black circle in the picture represents clinopyroxene, and the open circle represents orthopyroxene; A—diopside, B—subdiopside, C—enstatite, D—ordinary Pyroxene, E—clinopyroxene;
1 and 2 are the trend lines of granulite and igneous clinopyroxene coexisting with calcium-poor pyroxene respectively (Zhang Ruyuan et al., 1981)
(1) Orthopyroxene
The magnesia-iron ratio of orthopyroxene in granulite xenoliths XMg=MgO/(MgO+FeO) is 0.41 to 0.50, most of which are close to 0.5. In the MgO and FeO* content map of Opx (Figure 4-25A), orthopyroxene falls in the granulite of the North China terrane (Zhang Ruyuan et al., 1981; Shen Qihan et al., 1992; He Gaopin et al., 1991; Yan Yuehua et al., 1988; Jin Shiqin et al., 1986; Wang Shiqi et al., 1994) and Hannuoba xenolith granulite clinopyroxene (Chen Shaohai et al., 1998) are richer in Mg and poorer in Fe than the former, but richer and poorer in Fe than the latter. Mg. Relevant studies have shown that the CaO content of orthopyroxene in metamorphic rocks does not exceed 1.5%, and the increase in CaO content reflects the increase in the temperature of its formation. Experimental research on the MgO-Al2O3-SiO2 system shows that the Al2O3 content of orthopyroxene increases with increasing temperature (Boyd, 1973). In orthopyroxene that is not intergrown with garnet, the Al2O3 content increases with pressure. (Obata, 1976). The CaO content of orthopyroxene in granulite xenoliths in this area is all less than 1.5%, ranging from 0.49% to 1.31%, with an average of 0.98%; the Al2O3 content is 0.47% to 1.81%, most of which are less than 1.5%. Compared with orthopyroxene in Precambrian granulite in North China, it shows the characteristics of high CaO and low Al2O3 and Na2O (Table 4-18), reflecting its formation temperature is higher and the pressure is slightly lower.
Figure 4-25 MgO-FeO* variation diagram of pyroxene and biotite in granulite xenoliths
See the text for the source of the data in the figure Narrative
Table 4-18 Comparison of the average composition (wB/%) of pyroxene minerals in granulite xenoliths and granulite in the Precambrian terrane of the North China Craton
Note: 1) The data are quoted from Zhang Ruyuan et al. (1981), Shen Qihan et al. (1992), He Gaopin et al. (1991); 2) Quoted from Yan Yuehua et al. (1988); 3) Quoted from Jin Shiqin et al. (1986); 4) Quoted from Wang Shiqi et al. (1994).
Perilla pyroxene begins to form at a higher temperature in metamorphic rocks, and it is generally believed that it marks the beginning of granulite phase metamorphism. However, due to granulite trappingThe original rock of the body may be gabbro-type rocks, so it is necessary to determine whether these orthopyroxenes originated from early magma or were formed by later granulite phase metamorphism. Dobretsov (1970) once proposed the following discriminant formula to determine the origin of orthopyroxene (the element symbol in the formula represents the corresponding cation number in the structural formula of pyroxene, and the Fe3+ correction adopts the method of Zheng Qiaorong (1983)):
D(x)=-13.5+59.6AlIV+16.6Fe3++21.2Fe2++15.9Mn-5.12Mg+0.9Na
If D(x) is greater than 0, it is granulite phase (no pomegranate pyroxene (sub-stone); if D(x) is less than 0, it is a pyroxene of gabbro type derived from magma. The xenolith granulite clinopyroxene D(x) in this area are all greater than 0, indicating the origin of granulite phase metamorphism. In addition, Bhattacharyya (1971) once distinguished clinopyroxene from igneous and metamorphic origins based on the content of Al2O3 and (MgO+FeO+Fe2O3), and believed that those with (MgO+FeO+Fe2O3) + 0.775 and Al2O3 greater than 44.304 were of metamorphic origin. On the contrary, it is caused by magma. The values of orthopyroxene in this area are all greater than 44.304, further indicating that they are the products of granulite phase metamorphism.
According to the mathematical statistics of Dobretsov et al. (1973), the formation conditions of orthopyroxene are comprehensively reflected in its composition, and the following discriminant is proposed:
D (x) = - 4282+683Si+2192AlⅥ+2182Fe3++1455Mn+1442Mg+1427Ca+1700 (Na+K)
It is believed that D(x) is greater than 0, which belongs to the high-temperature granulite phase; D(x) is less than 0, It belongs to hornblende granulite facies. The D(x) values of clinopyroxene in the granulite xenoliths in this area are all greater than 0, indicating that the formation temperature is relatively high and has reached the high-temperature granulite phase.
(2) Clinopyroxene
Clinopyroxene in granulite xenoliths (Table 4-18) shows rich Mg, poor Fe, low Al, and Ti , Na's characteristics. The clinopyroxene-magnesium-iron ratio XMg is 0.5 to 0.64. In the MgO-FeO* diagram of Cpx (Figure 4-25B), its MgO and FeO* contents are obviously different from the clinopyroxene in the granulite of the Precambrian terrane in North China, and different from the granulite in the Hannuoba basalt in Hebei. Grainite xenoliths are similar to clinopyroxene (Fan Qicheng et al., 1996). Experimental data show that the AlVI and Na contents of clinopyroxene generally increase with increasing pressure (Zhang Ruyuan et al., 1981; Wood, 1974); the CaO content can reflect the temperature change trend and increases with decreasing temperature, and pressure has an effect on it. The effect is negligible below 1000°C (MacGregor et al., 1976); the effect of Ti is opposite to that of Na. As the pressure increases, the content of titanium pyroxene molecules decreases. When the pressure is greater than 10kbar, they do not appear (Wood, 1974). Granulite xenoliths in this areaThe Al content of clinopyroxene is low, among which AlVI is 0.028-0.103, with an average of 0.074, which is greater than the average content of AlIV, 0.019, and there is no obvious correlation between the two. CaO content is 17% ~ 22.4%. The content of Na varies greatly, from 0.003 to 0.04, with an average of 0.021. Na and AlVI are negatively correlated. The content of Ti is very low, with an average of 0.007. Compared with the clinopyroxene in the granulites of the North China terrane, the CaO content of clinopyroxene in this area is slightly higher and the AlVI and Na contents are low (Table 4-19), indicating that the temperature at which they were formed was slightly higher and the pressure was lower, but The Ti content is similar, which may mean that the pressure is not too low. Calculating the end-member molecules based on mineral chemical composition, the jadeite molecules (NaAl [Si2O6]) of clinopyroxene are very low, with a maximum of no more than 2%, which means that its formation pressure will not be too high. In short, the chemical composition characteristics of clinopyroxene in granulite xenoliths in this area indicate that its formation temperature is relatively high and the pressure is relatively low, which is consistent with the results of the composition of clinopyroxene.
The composition of clinopyroxene is closely related to the origin of the host rock. Vejnar (1975) used Ti-Al2O3 and (Si+Al)-Al2O3 variation diagrams to divide clinopyroxene into two categories: metamorphic origin and magmatic origin. In both diagrams, clinopyroxene in this area falls into the metamorphic origin zone (Figure 4-26). Calculated based on the method proposed by Dobretsov (1970) to identify the genetic type of clinopyroxene, the clinopyroxene in the granulite xenoliths in this area is also of granulite phase metamorphic origin.
Figure 4-26 Ti-Al2O3 and (Si+Al)-Al2O3 diagrams of granulite xenolith clinopyroxene
A— Metamorphic; B-magmatic origin, the black dots represent the clinopyroxene in the granulite, and the shaded area represents the clinopyroxene composition area in the cumulate rock
As can be seen from Figure 4-24, Clinopyroxene is basically distributed along the granulite phase clinopyroxene trend line determined by Binns (Zhang Ruyuan et al., 1981). The line connecting the projection points of symbiotic lherzolite in the figure intersects with Wo-En between Wo70En30 and Wo80En20, indicating that the symbiotic lherrite is in equilibrium symbiosis (Yan Yuehua, 1997), and is similar to the equilibrium symbiosis of lher pyroxene in igneous rocks (Deer, 1978).
2. Biotite
Granulite xenoliths generally contain biotite, with content ranging from 1% to 10%. Most of them are brown-red (Ng). This is A sign of high TiO2 content. Ti has a strong influence on the color of mica, especially biotite. Generally, the higher the TiO2 content, the redder the mica color is (Yan Yuehua et al., 1988).
The chemical composition of biotite is shown in Table 4-17. It can be seen from the table that its composition is relatively rich in Mg, and most of the Mg/Fe ratios are greater than 2, which already belongs to the category of phlogopite. However, there is no phlogopite in the granulites of the Precambrian terrane in North China, and the two types of granulites There are obvious differences in the mica composition of the rocks (Figure 4-25C). These biotites are poorer in magnesium and richer in iron than the typical phlogopite in Shandong kimberlites (Zhou Zuoxia, 1988) and the phlogopite xenoliths contained in the same host rock in this area, and are different from the basic granulites in Queensland, Australia. Biotite in xenoliths (Rudnick et al., 1987) is equivalent. In addition, the biotite in the granulite xenoliths in this area also shows the characteristics of high Ti, low Mn, and low AlVI and Si. The TiO2 content is 2.4% to 5.2%, MnO is 0 to 0.1% (mostly less than 0.05%), Al2O3 is 13.8% to 15.2%, AlVI is 0 to 0.19 in the chemical formula calculated based on 22 O, and AlIV is relatively High, AlⅣ/Si=0.419～0.469, there is no obvious correlation between AlⅣ and AlⅥ.
The above-mentioned characteristics of biotite with high Ti, low Mn, and low AlVI clearly indicate that the xenolith metamorphism has reached the granulite phase. High Ti, low Mn, and low AlVI are the characteristics of granulite phase biotite (Yan Yuehua et al., 1988). Many researchers have proven that the Ti content in biotite is a function of the degree of metamorphism, and the Ti content increases with the degree of metamorphism (Zhang Ruyuan et al., 1981). Ti in amphibolite phase biotite is generally less than 0.3 (calculated as 22 O), while Ti in granulite phase biotite is higher (Guitdotti, 1984). The biotite Ti in granulite xenoliths in this area is generally greater than 0.3, with the highest reaching 0.564. The AlVI of biotite is also a function of the degree of metamorphism. The AlVI of biotite in low-grade metamorphic rocks is high, and the AlVI of biotite is low when the degree of metamorphism is high. The AlVI of granulite phase biotite is usually less than 0.55. The AlVI of biotite in non-argillaceous granulite phase rocks is lower (Guitdotti, 1984). The highest biotite AlVI in this area is only 0.19, which is much less than 0.55. In addition, in the TiO2-100Fe/(Fe+Mg) diagram of metamorphic biotite (Zhang Ruyuan et al., 1983), the composition points of the biotite samples all fall in the granulite phase area (Figure 4-27). Therefore, it can be said that the biotite in the xenolith granulite in this area has the characteristics of granulite phase and is the product of granulite phase metamorphism.
Figure 4-27 Illustration of biotite TiO2-100Fe/(Fe+Mg) in granulite xenoliths
3. Plagioclase< /p>
Plagioclase is the main light-colored rock-forming mineral in granulite xenoliths. Generally, plagioclase is sensitive to the chemical composition, intergenetic minerals and metamorphism grade of the host rock. The An content of plagioclase formed during regional metamorphism increases with the increase of temperature. The amphibolite phase plagioclase The composition is mainly in the range of labradorite or andesine, while the granulite phase can reach the range of peclase and labradorite (Ye Huiwen, 1986). According to the chemical composition of plagioclase, its end-member chemical composition is calculated. The plagioclase An in the granulite xenoliths in this area is 42 to 74, which are andesine, labradorite and peclase respectively.
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